Encyclopedia of the World's Coastal Landforms

Encyclopedia of the World‘s Coastal Landforms

Eric C. F. Bird Editor

Encyclopedia of the World‘s Coastal Landforms Volume I

With 1379 Figures

123

Editor Eric C.F. Bird, M.Sc. (Lond. and Melb.) Ph.D. (ANU) Principal Fellow in Geomorphology University of Melbourne and Director, Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected]

ISBN 978-1-4020-8638-0 This publication is available also as: Electronic publication under ISBN 978-1-4020-8639-7 Print and electronic bundle under ISBN 978-1-4020-8640-3 DOI 10.1007/978-1-4020-8640-3 Springer Dordrecht Heidelberg London New York Library of Congress Control Number: 2009938922 © Springer Science+Business Media B.V. 2010 All rights reserved: Chapters 8.2, 8.15, 9.4, 10.1, 11.2, 11.3, 18.1 and 20.5 Every effort has been made to contact the copyright holders of the figures and tables which have been reproduced from other sources. Anyone who has not been properly credited is requested to contact the publishers, so that due acknowledgment may be made in subsequent editions. No part of this work may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, microfilming, recording or otherwise, without written permission from the Publisher, with the exception of any material supplied specifically for the purpose of being entered and executed on a computer system, for exclusive use by the purchaser of the work. Cover figures: Images 5084487, 5078544 and 5333716 from photos.com. © 2010 photos.com Printed on acid-free paper Springer is part of Springer Science+Business Media (www.springer.com)

Preface Encyclopedia of the World’s Coastal Landforms is an account of the various kinds of coastal morphology found on the world’s coastlines. It is presented in 224 chapters, based on national, state, or county divisions, and arranged in a sequence that proceeds counterclockwise around the continents. The sequence begins with Alaska, proceeds around North America to Arctic Canada, then Greenland and finally, Iceland. Another counterclockwise sequence begins with Norway and proceeds by way of Europe and the Mediterranean around Africa to India, Southeast Asia, China, Korea, and the Pacific and Arctic coasts of Russia. Interruptions are made at appropriate points to include the Great Lakes, the Caspian Sea, and islands such as Britain, Madagascar, and Japan. The sequence is completed with the Philippines, Indonesia, Papua New Guinea, Australia, New Zealand, the islands of the Pacific, Atlantic, and Indian Oceans, and finally Antarctica. This work was preceded by The World’s Coastline, edited by Eric C. F. Bird and Maurice L. Schwartz and published by Van Nostrand Reinhold, New York, in 1985. An internet version was published by Kluwer Academic Publishers in 2004. This 2010 edition has also been produced with the aid of 123 contributors. It includes 1,379 illustrations and cites numerous references. The focus is on coastal geomorphology, with accounts of features seen along the coastline, illustrating sites of special scientific interest. It provides access to the large amount of data that has been assembled, to ensure that it will not just disappear over the next few years. It is now much more difficult for scientists to make broad coastal surveys as they are required to specialise more, which usually means they work on small selected coastal sites. This book will allow the scientists to place their work in a global context and make comparisons with similar coasts in other parts of the world. Coastal geomorphology has several interest groups. Some focus on the shaping of coastal landforms in relation to geology, processes, variations in climate, and the relative levels of land and sea. Others concentrate on coastline changes measured over specified periods, with analyses of their causes. Others are interested in nearshore processes and responses, particularly beaches, where morphodynamic models have been most successful. Others look to the coastal area for evidence of Quaternary history, notably changes in land- and sea level and climatic variations. Others examine the sources and patterns of movement of coastal sediment, or the array of physical-, chemical-, and biological weathering processes in the coastal zone. Coastal managers want information on how landforms are changing, what processes are at work on them, how they will change in the future, and what effects will result from the development of various kinds. The chapters presented in this version of Encyclopedia of the World’s Coastal Landforms reflect the interests of their authors, but an attempt is made to serve the wide range of interests of coastal geomorphologists. The work was designed initially for scientists needing to identify locations and distributions of coastal features and to make comparisons between coastal landforms in various parts of the world. However, it will also be of value to people seeking general information on coastal landforms in the various countries around the world, including those planning travel and exploration of coastal regions. Armchair travellers can use it to make substantial journeys around the world’s coasts, enjoying a succession of visual images. Indeed, it is unique in presenting the wide variations in scenery and environment that occur on the world’s coastline. I would like to thank Lew Ward for processing the edited material and illustrations. In assembling the work, I had help from librarians in the University of Melbourne, notably David Jones, Map Curator in the Education Resource Centre. In Black Rock I had assistance from Dr. Juliet Bird, and from Myriam, Audrey, Ingrid, and Céline Vinot. Eric C. F. Bird Geostudies, Melbourne, Australia January 2010

Contents Volume I Editor .......................................................................................................................................................... xxiii List of Contributors ...................................................................................................................................... xxv 1.0

USA – Editorial Introduction ........................................................................................................................... 1 Eric Bird

1.1

Alaska ................................................................................................................................................................ 5 Jesse Walker · Molly McGraw

1.2

Washington ..................................................................................................................................................... 15 Maurice Schwartz · Thomas Terich

1.3

Oregon ............................................................................................................................................................ 33 Paul D. Komar

1.4

California ........................................................................................................................................................ 43 Eric Bird

1.5

Texas ............................................................................................................................................................... 53 Robert Morton

1.6

Louisiana ........................................................................................................................................................ 61 Ioannis Georgiou · Mark Kulp · Michael Miner · Duncan FitzGerald

1.7

Mississippi ...................................................................................................................................................... 69 Ervin Otvos

1.8

Alabama .......................................................................................................................................................... 77 Ervin Otvos

1.9

Florida ............................................................................................................................................................ 83 Charles Finkl

1.10

Georgia ........................................................................................................................................................... 89 Miles Hayes

1.11

South Carolina ............................................................................................................................................... 93 Miles Hayes

1.12

North Carolina ............................................................................................................................................... 99 Stanley R. Riggs

1.13

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey) ........................................ 107 Eric Bird

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Contents

1.14

New York and New England ......................................................................................................................... 113 Henry Bokuniewicz

1.15

Great Lakes (USA) ........................................................................................................................................ 121 Mary-Louise Byrne

1.16

Hawaii ........................................................................................................................................................... 125 Charles H. Fletcher · Eden J. Feirstein

2.0

Canada – Editorial Introduction ................................................................................................................. 133 Eric Bird

2.1

British Columbia .......................................................................................................................................... 135 John J. Clague

2.2

New Brunswick and Nova Scotia ................................................................................................................. 141 Eric Bird

2.3

Gulf of St. Lawrence ..................................................................................................................................... 155 Eric Bird

2.4

Saint-Pierre-et-Miquelon Islands ................................................................................................................ 163 Jean-Marie M. Dubois

2.5

Newfoundland .............................................................................................................................................. 167 Eric Bird

2.6

Northern Canada ......................................................................................................................................... 169 Wayne Pollard

2.7

Great Lakes (Canada) ................................................................................................................................... 177 Mary-Louise Byrne

3.0

Pacific Central America – Editorial Introduction ....................................................................................... 183 Eric Bird

3.1

Mexico, Pacific Coast ................................................................................................................................... 185 Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez

3.2

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama ...................... 187 Anja Scheffers · Tony Browne

3.3

Pacific Coast of Colombia ............................................................................................................................ 193 Iván Correa · Robert Morton

4.0

South America – Editorial Introduction ..................................................................................................... 199 Eric Bird

4.1

Ecuador ......................................................................................................................................................... 201 H. Ayon · W. Jara

Contents

4.2

Peru ............................................................................................................................................................... 207 Vidal Taype Ramos · Eric Bird

4.3

Chile .............................................................................................................................................................. 213 José Araya-Vergara

4.4

Argentina ...................................................................................................................................................... 219 Enrique Schnack · Jorge Pousa · Germán Bértola · Federico Isla

4.5

Uruguay ........................................................................................................................................................ 227 Eric Bird

4.6

Brazil ............................................................................................................................................................. 231 Dieter Muehe

4.7

French Guiana .............................................................................................................................................. 239 J. Turenne

4.8

Surinam ........................................................................................................................................................ 243 Norbert Psuty

4.9

Guyana .......................................................................................................................................................... 245 Eric Bird

4.10

Venezuela ...................................................................................................................................................... 249 L. Ellenberg

4.11

Caribbean Coast of Colombia ..................................................................................................................... 259 Iván Correa · Robert Morton

4.12

Caribbean Coasts, Panama to Belize ............................................................................................................ 265 Anja Scheffers · Tony Browne

4.13

Mexico, Caribbean Coast ............................................................................................................................. 269 Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez

5.0

Caribbean Islands – Editorial Introduction ................................................................................................ 271 Eric Bird

5.1

Cuba ............................................................................................................................................................. 273 Ridel Rodríguez

5.2

Jamaica ......................................................................................................................................................... 279 John Norrman · Tommy Lindell

5.3

Hispaniola (Haiti and the Dominican Republic) ........................................................................................ 285 Anja Scheffers · Tony Browne

5.4

Puerto Rico ................................................................................................................................................... 289 Jack Morelock ∙ Wilson Ramirez ∙ Maritza Barreto

ix

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Contents

5.4.1

Virgin Islands ....................................................................................................................................... 295 Maurice Schwartz

5.5

Lesser Antilles .............................................................................................................................................. 299 Gillian Cambers

6.0

Atlantic Ocean Islands – Editorial Introduction ........................................................................................ 311 Eric Bird

6.1

Bahamas ....................................................................................................................................................... 313 Alan K. Craig

6.2

Bermuda ....................................................................................................................................................... 317 Eric Bird

6.3

Greenland ..................................................................................................................................................... 319 Eric Bird

6.4

Iceland .......................................................................................................................................................... 323 Eggert Larusson

6.5

Other Atlantic Ocean Islands ...................................................................................................................... 333 Jim Hansom

7.0

British Isles – Editorial Introduction .......................................................................................................... 345 Eric Bird

7.1

England and Wales ....................................................................................................................................... 347 Eric Bird

7.2

Cumbria ....................................................................................................................................................... 353 Eric Bird

7.3

Lancashire, Merseyside and Cheshire ......................................................................................................... 358 Eric Bird

7.4

The Isle of Man .............................................................................................................................................. 361 Eric Bird

7.5

North Wales and Anglesey ........................................................................................................................... 366 Eric Bird

7.6

West Wales .................................................................................................................................................... 369 Eric Bird

7.7

South Wales ................................................................................................................................................... 375 Eric Bird

7.7.1

The Severn Estuary ............................................................................................................................... 384 Eric Bird

Contents

7.8

Gloucestershire, Somerset and North Devon .............................................................................................. 386 Eric Bird

7.9

Cornwall

7.9.1

The North Coast of Cornwall .............................................................................................................. 394 Eric Bird

7.9.2

The South Coast of Cornwall ............................................................................................................... 401 Eric Bird

7.10

Isles of Scilly ................................................................................................................................................. 410 Eric Bird

7.11

South Devon.................................................................................................................................................. 413 Eric Bird

7.12

Dorset ........................................................................................................................................................... 420 Eric Bird

7.13

Hampshire ..................................................................................................................................................... 429 Eric Bird

7.14

Isle of Wight ................................................................................................................................................. 433 Eric Bird

7.15

Sussex ........................................................................................................................................................... 438 Eric Bird

7.16

Kent .............................................................................................................................................................. 443 Eric Bird

7.17

Essex .............................................................................................................................................................. 447 Eric Bird

7.18

Suffolk ........................................................................................................................................................... 450 Eric Bird

7.19

Norfolk .......................................................................................................................................................... 454 Eric Bird

7.20

Lincolnshire ................................................................................................................................................. 459 Eric Bird

7.21

Yorkshire ...................................................................................................................................................... 461 Eric Bird

7.22

Durham, Tyne and Wear............................................................................................................................... 466 Eric Bird

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7.23

Northumberland ........................................................................................................................................... 468 Eric Bird

7.24

Scotland ........................................................................................................................................................ 473 William Ritchie · Alastair Dawson

7.24.1

South-East Scotland.............................................................................................................................. 477 William Ritchie

7.24.2

Fife......................................................................................................................................................... 481 William Ritchie

7.24.3

Angus and Aberdeenshire .................................................................................................................... 485 William Ritchie

7.24.4

Fraserburgh to Moray Firth and Duncansby Head ............................................................................. 489 William Ritchie

7.24.5

The Orkney and Shetland Islands......................................................................................................... 493 Alastair Dawson

7.24.6

North and North West Scotland (Duncansby Head to Cape Wrath and Fort William) ..................... 499 Alastair Dawson

7.24.7

The Inner Hebrides ............................................................................................................................... 505 Alastair Dawson

7.24.8

The Outer Hebrides .............................................................................................................................. 515 Alastair Dawson

7.24.9

Fort William to the Clyde .................................................................................................................... 523 Alastair Dawson

7.24.10

South-West Scotland............................................................................................................................. 529 William Ritchie

7.25

Northern Ireland .......................................................................................................................................... 535 Andrew Cooper

7.26

Ireland .......................................................................................................................................................... 545 Eric Bird

7.27

Channel Islands ............................................................................................................................................ 557 Eric Bird

8.0

Europe – Editorial Introduction .................................................................................................................. 569 Eric Bird

8.1

Norway ......................................................................................................................................................... 571 Tormod Klemsdal

Contents

8.1.1

Svalbard and Jan Mayen ....................................................................................................................... 581 Tormod Klemsdal

8.2

Sweden .......................................................................................................................................................... 589 John Norrman

8.3

Finland ......................................................................................................................................................... 595 Jouko Alestalo · Olavi Granö

8.3.1

Russian Gulf of Finland ........................................................................................................................ 601 Alexey Porotov

8.4

Estonia .......................................................................................................................................................... 605 Kaarel Orviku · Are Kont · Hannes Tõnisson

8.5

Latvia ............................................................................................................................................................ 613 Guntis Eberhards · Viktor Brenners

8.6

Lithuania ...................................................................................................................................................... 619 Eric Bird

8.7

Kaliningrad .................................................................................................................................................. 623 Alexey Porotov

8.8

Poland ........................................................................................................................................................... 627 Karol Rotnicki · Joanna Rotnicka

8.9

Germany ....................................................................................................................................................... 639 Klaus Schwarzer · Horst Sterr

8.10

Denmark ....................................................................................................................................................... 649 Troels Aagaard

8.11

The Netherlands ........................................................................................................................................... 665 Eric Bird

8.12

Belgium ........................................................................................................................................................ 669 Guy De Moor · André Ozer · Irénée Heyse

8.13

France

8.13.1

North Coast of France .......................................................................................................................... 673 Eric Bird

8.13.2

West Coast of France ........................................................................................................................... 683 Eric Bird

8.13.3

Mediterranean France .......................................................................................................................... 691 Eric Bird

8.13.4

Corsica ................................................................................................................................................. 695 Eric Bird

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Contents

8.14

8.14.1

Spain ............................................................................................................................................................. 699 Eric Bird Gibraltar ............................................................................................................................................... 711 Eric Bird

8.15

Portugal ........................................................................................................................................................ 713 Carlos Morais

8.16

Italy ............................................................................................................................................................... 719 Paolo Ciavola

8.16.1

Sardinia ................................................................................................................................................ 731 Eric Bird

8.16.2

Sicily ..................................................................................................................................................... 741 Paolo Ciavola

8.17

Malta ............................................................................................................................................................. 751 George Said · John Schembri

8.18

Slovenia ........................................................................................................................................................ 761 Roger Charlier

8.19

Croatia .......................................................................................................................................................... 763 Roger Charlier

8.20

Bosnia-Herzegovina ..................................................................................................................................... 769 Roger Charlier

8.21

Montenegro .................................................................................................................................................. 771 Roger Charlier

8.22

Albania ......................................................................................................................................................... 773 Yurii Shuisky

8.23

Greece ........................................................................................................................................................... 775 Eric Bird

8.24

Crete ............................................................................................................................................................. 783 Anja Scheffers · Tony Browne

8.25

Bulgaria ........................................................................................................................................................ 785 Eric Bird

8.26

Romania ....................................................................................................................................................... 789 Roger Charlier · Constance Charlier

8.27

Ukraine ......................................................................................................................................................... 795 Yurii Shuisky

Contents

8.28

Sea of Azov ................................................................................................................................................... 799 Eric Bird

8.29

Republic of Georgia ..................................................................................................................................... 805 Pavel Kaplin · Andrei Selivanov Index Maps ................................................................................................................................................. 1471 List of Entries ............................................................................................................................................. 1483 Subject Index .............................................................................................................................................. 1489 Volume II Editor .......................................................................................................................................................... xxiii List of Contributors ...................................................................................................................................... xxv

9.0

Russian Federation – Editorial Introduction .............................................................................................. 809 Eric Bird

9.1

Russian Gulf of Finland ............................................................................................................................... 810 Alexey Porotov

9.2

Russian Black Sea Coast ............................................................................................................................... 811 Pavel Kaplin · Andrei Selivanov

9.3

The Pacific Coast of Russia .......................................................................................................................... 815 Pavel Kaplin · Andrei Selivanov

9.4

The Arctic Coast of Russia ........................................................................................................................... 823 Andrei Selivanov

9.5

Russian Caspian Coast.................................................................................................................................. 829 Svetlana Lukyanova

9.6

Kaliningrad ................................................................................................................................................... 829 Alexey Porotov

10.0

Middle East – Editorial Introduction .......................................................................................................... 831 Eric Bird

10.1

Turkey ........................................................................................................................................................... 833 Oguz Erol

10.2

Cyprus .......................................................................................................................................................... 841 Yaacov Nir

10.3

Syria ............................................................................................................................................................... 849 Eric Bird

10.4

Lebanon ........................................................................................................................................................ 851 Eric Bird

xv

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Contents

10.5

Israel (with the Gaza Strip) ........................................................................................................................... 853 Yaacov Nir

11.0

Caspian and Aral Seas – Editorial Introduction ......................................................................................... 859 Eric Bird

11.1

Iran – Caspian Sea Coast .............................................................................................................................. 861 Eric Bird

11.2

Turkmenistan ................................................................................................................................................ 867 Andrei Selivanov

11.3

Kazakhstan ................................................................................................................................................... 871 Andrei Selivanov

11.4

Russian Caspian Coast ................................................................................................................................. 875 Svetlana Lukyanova

11.5

Azerbaijan .................................................................................................................................................... 885 Pavel Kaplin · Andrei Selivanov · Svetlana Lukyanova

11.6

Aral Sea ......................................................................................................................................................... 889 Eric Bird

12.0

North Africa – Editorial Introduction ........................................................................................................ 891 Eric Bird

12.1

Egypt (Mediterranean) ................................................................................................................................ 893 Anja Scheffers ∙ Tony Browne

12.2

Libya ............................................................................................................................................................. 897 Maurice Schwartz

12.3

Tunisia .......................................................................................................................................................... 899 Eric Bird

12.4

Algeria .......................................................................................................................................................... 903 M. Larid

12.5

Morocco ........................................................................................................................................................ 907 Anja Scheffers

13.0

West Africa – Editorial Introduction .......................................................................................................... 915 Eric Bird

13.1

Mauritania .................................................................................................................................................... 917 Don Vermeer

13.2

Senegal and Gambia ..................................................................................................................................... 921 André Guilcher

Contents

13.3

Guinea Bissau................................................................................................................................................ 927 E.S. Diop

13.4

Republic of Guinea ....................................................................................................................................... 929 E.S. Diop

13.5

Sierra Leone ................................................................................................................................................. 931 Eric Bird

13.6

Liberia .......................................................................................................................................................... 935 Eric Bird

13.7

Ivory Coast ................................................................................................................................................... 939 Anja Scheffers

13.8

Ghana ........................................................................................................................................................... 943 L.A. Dei

13.9

Togo and Benin ............................................................................................................................................ 947 André Guilcher

13.10 Nigeria .......................................................................................................................................................... 949 Etop Usoro 13.11 Cameroon and Equatorial Guinea ............................................................................................................... 953 Maurice Schwartz 13.12 Gabon, Congo, Cabinda and Zaïre .............................................................................................................. 957 P. Giresse 13.13 Angola .......................................................................................................................................................... 963 André Guilcher 13.14 Namibia ........................................................................................................................................................ 969 Eric Bird · A. Goudie · H. Viles

14.0

South Africa – Introduction ........................................................................................................................ 975 Gerald G. Garland

14.1

South Africa ................................................................................................................................................. 979 Andrew Cooper

15.0

East Africa – Editorial Introduction ........................................................................................................... 987 Eric Bird

15.1

Mozambique ................................................................................................................................................. 989 Maria Eugénia Soares de Albergaria Moreira

15.2

Tanzania ....................................................................................................................................................... 995 Eric Bird

xvii

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Contents

15.3

Kenya .......................................................................................................................................................... 1003 Francis Ojany

15.4

Somalia and Djibouti ................................................................................................................................. 1007 Eric Bird

15.5

Eritrea ......................................................................................................................................................... 1011 Eric Bird

15.6

Sudan .......................................................................................................................................................... 1013 Eric Bird

15.7

Egypt, Red Sea Coast ................................................................................................................................. 1015 Anja Scheffers · Tony Browne

15.8

Jordan ......................................................................................................................................................... 1019 Yaacov Nir

16.0

South-West Asia – Editorial Introduction ................................................................................................ 1021 Eric Bird

16.1

Saudi Arabia, Red Sea Coast ...................................................................................................................... 1023 Eric Bird

16.2

Yemen ......................................................................................................................................................... 1027 Eric Bird

16.3

Southern Arabia and Oman ....................................................................................................................... 1029 Eric Bird

16.4

United Arab Emirates ................................................................................................................................ 1033 Eric Bird

16.5

Qatar ........................................................................................................................................................... 1037 Eric Bird

16.6

Bahrein ....................................................................................................................................................... 1041 Eric Bird

16.7

Saudi Arabia, Persian Gulf Coast .............................................................................................................. 1045 Eric Bird

16.8

Kuwait and Iraq .......................................................................................................................................... 1047 Eric Bird

16.9

Iran ............................................................................................................................................................. 1051 Rodman E. Snead

17.0

South Asia – Editorial Introduction .......................................................................................................... 1057 Eric Bird

Contents

17.1

Pakistan ...................................................................................................................................................... 1059 Rodman E. Snead

17.2

India ........................................................................................................................................................... 1065 G. N. Nayak · P. T. Hanamgond

17.3

Sri Lanka .................................................................................................................................................... 1071 Eric Bird

17.4

Bangladesh ................................................................................................................................................. 1077 Rodman E. Snead

17.5

Burma (Myanmar) ..................................................................................................................................... 1081 Kyaw Saw Lynn

18.0

Indian Ocean Islands – Editorial Introduction ......................................................................................... 1087 Eric Bird

18.1

Madagascar ................................................................................................................................................ 1089 Jean-Michel Lebigre

18.2

Indian Ocean Islands ................................................................................................................................. 1097 Wong Poh Poh

19.0

South-East Asia – Editorial Introduction ................................................................................................. 1111 Eric Bird

19.1

Thailand Andaman Sea Coast .................................................................................................................... 1113 Sanit Aksornkoae · Eric Bird

19.2

Malaysia – Introduction ............................................................................................................................. 1117 Teh Tiong Sa · Yap Hui Boon

19.3

Singapore .................................................................................................................................................... 1129 Wong Poh Poh

19.3.1

Thailand: Gulf of Thailand Coast ...................................................................................................... 1135 Sanit Aksornkoae · Eric Bird

19.4

Brunei (Negara Brunei Darussalam) ......................................................................................................... 1141 Gabriel Yong

19.5

Cambodia ................................................................................................................................................... 1145 Eric Bird

19.6

Vietnam ...................................................................................................................................................... 1147 D. Eisma

19.7

Philippines ................................................................................................................................................. 1151 Eric Bird

xix

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Contents

19.8

Indonesia .................................................................................................................................................... 1157 Otto Ongkosongo

19.9

East Timor .................................................................................................................................................. 1171 Wong Poh Poh

19.10 Papua New Guinea ..................................................................................................................................... 1175 Eric Bird

20.0

East Asia – Editorial Introduction ............................................................................................................ 1187 Eric Bird

20.1

China .......................................................................................................................................................... 1189 Chen Jiyu · Liu Cangzi · Dai Zhijun · Yu Zhiying

20.2

Taiwan ........................................................................................................................................................ 1197 John R. C. Hsu

20.3

North Korea ................................................................................................................................................ 1201 D. Eisma

20.4

South Korea ................................................................................................................................................ 1207 D. Eisma

20.5

Japan ........................................................................................................................................................... 1213 Kazuyuki Koike

21.0

Australia – Editorial Introduction ............................................................................................................ 1225 Eric Bird

21.1

New South Wales ........................................................................................................................................ 1229 Bruce Thom

21.1.1

Lord Howe Island – (New South Wales) ............................................................................................ 1239 Eric Bird

21.1.2

Norfolk Island .................................................................................................................................... 1247 Eric Bird

21.2

Queensland ................................................................................................................................................. 1255 David Hopley · Scott Smithers

21.3

Northern Territory ..................................................................................................................................... 1267 Eric Bird

21.4

Western Australia ....................................................................................................................................... 1279 Ian Eliot · Eric Bird

21.5

South Australia .......................................................................................................................................... 1293 C. R. Twidale · J. A. Bourne

Contents

21.6

Victoria Introduction ................................................................................................................................. 1305 Eric Bird

21.6.1

Victoria: The Western District (Discovery Bay to Mepunga) ............................................................ 1311 Eric Bird

21.6.2

Victoria: The Port Campbell Coast (Mepunga to Princetown) ......................................................... 1319 Eric Bird

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale) .......................... 1325 Eric Bird

21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)............................................................ 1337 Eric Bird

21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head) .................................................... 1349 Eric Bird

21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island.................................................... 1359 Eric Bird

21.6.7

South Gippsland, Wilsons Promontory and Corner Inlet ................................................................ 1369 Eric Bird

21.6.8

Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe) ................................................... 1377 Eric Bird

21.6.9

Victoria: The Gippsland Lakes ........................................................................................................... 1385 Eric Bird

21.7

Tasmania .................................................................................................................................................... 1395 Joanna Ellison

22.1

New Zealand ............................................................................................................................................... 1403 Terry Healy

22.2

The Chatham Islands ................................................................................................................................. 1413 Eric Bird

23.0

Pacific Ocean Islands – Editorial Introduction ......................................................................................... 1423 Eric Bird

23.1

New Caledonia and the Loyalty Islands ..................................................................................................... 1425 Jean-Michel Lebigre · Eric Bird

23.2

Fiji ............................................................................................................................................................... 1437 Patrick Nunn

23.3

Society Islands ............................................................................................................................................ 1451 André Guilcher

xxi

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Contents

23.4

Other Pacific Islands .................................................................................................................................. 1455 Joanna Ellison

24.1

Antarctic Coast .......................................................................................................................................... 1463 Eric Bird Index Maps ................................................................................................................................................. 1471 List of Entries ............................................................................................................................................. 1483 Subject Index .............................................................................................................................................. 1489

Editor Eric C.F. Bird, M.Sc. (Lond. and Melb.) Ph.D. (ANU) Principal Fellow in Geomorphology University of Melbourne and Director, Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected]

List of Contributors Troels Aagaard Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen Denmark [email protected] Sanit Aksornkoae Thailand Environment Institute (TEI) 16/151 Muang Thong Thani Bond Street Bangpood, Pakkred Nonthaburi 11120 Thailand Jouko Alestalo Department of Geography University of Turku Turku 20014 Finland José Araya-Vergara Department of Geography University of Chile Av. Portugal 84, Cas. 3387 Santiago de Chile Chile [email protected] H. Ayon Escuela Superior Politecnica del Litoral (ESPOL) Centro Nacional de Recursos Costeros Malecon No 100 y Loja Guayaquil Ecuador

Germán Bértola Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Centro de Geología de Costas y del Cuaternario Funes 3350, Mar del Plata Argentina [email protected] Eric C.F. Bird Geostudies 343 Beach Road Black Rock, Victoria 3193 Australia [email protected] Henry Bokuniewicz Marine Sciences Research Center State University of New York Stony Brook, NY 11794-5000 USA [email protected] Yap Hui Boon Hwa Chong Institution 661, Bukit Timah Road, 269734 Singapore J.A. Bourne School of Earth and Environmental Sciences Geology and Geophysics University of Adelaide Adelaide 5005 South Australia Viktor Brenners 114 Princess Street Kew, Victoria Australia Tony Browne†

Maritza Barreto Marine Sciences Department University of Puerto Rico at Mayagüez P.O. Box 1127 Lajas, PR Puerto Rico [email protected]

Mary-Louise Byrne Geography and Environmental Studies Wilfrid Laurier University Waterloo, Ontario Canada N2L 3C5 [email protected]

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List of Contributors

Gillian Cambers Sea Grant College Program University of Puerto Rico at Mayagüez P.O. Box 9011 Puerto Rico 681 [email protected] Liu Cangzi State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China Constance Charlier C.E.O. Inc. Tucson, AZ USA Roger Charlier 2 avenue du Congo box 23 B-1050 Brussels Belgium [email protected] Paolo Ciavola Dipartimento di Scienze della Terra Università di Ferrara Corso Ercole I d‘Este, 32 44100 Ferrara Italy [email protected] John J. Clague Department of Earth Sciences Simon Fraser University 8888 University Drive Burnaby, British Columbia Canada V5A 1S6 [email protected] Andrew Cooper School of Environmental Science University of Ulster Cromore Road Coleraine BT52 1SA Northern Ireland UK [email protected]

Iván Correa Departamento de Geología Universidad Eafit P.O. Box 3300 Medellin, Colombia South America [email protected] Alan K. Craig Florida Atlantic University Boca Raton, FL 33431 USA Alastair Dawson AICSM University of Aberdeen Elphinstone Road Aberdeen AB24 3UF Scotland UK [email protected] L.A. Dei Department of Geography & Regional Planning University of Cape Coast Box UC 0180, University Post Office Cape Coast Ghana [email protected] E.S. Diop University CAD Dakar Senegal Jean-Marie M. Dubois Département de géomatique appliquée Université de Sherbrooke Sherbrooke, Québec Canada J1K 2R1 [email protected] Guntis Eberhards Faculty of Geography and Earth Sciences University of Latvia Raiņa bulv. 19 Rīga Latvia [email protected]

List of Contributors

D. Eisma Oude Molstraat 12a 2513 BB Den Haag The Netherlands Ian Eliot Faculty of Natural and Agricultural Sciences University of Western Australia [email protected] L. Ellenberg Humboldt-Universität zu Berlin Geographisches Institut Chausseestraße 86 D- 10099 Berlin [email protected] Joanna Ellison Department of Geography University of Tasmania Launceston Tasmania 7001 Australia [email protected] Oguz Erol Ankara Üniversitesi 4. Kat Koridoru Ankara Turkey 6090 [email protected] Eden J. Feirstein Department of Geology and Geophysics School of Ocean and Earth Science and Technology (SOEST) University of Hawaii 1680 East-West Rd Honolulu, HI 96822 USA Charles Finkl Department of Geosciences Florida Atlantic University Boca Raton, FL 33431 USA [email protected]

Duncan FitzGerald Department of Earth Sciences Boston University 675 Commonwealth Avenue Boston, MA 02215 USA [email protected] Charles H. Fletcher Department of Geology and Geophysics School of Ocean and Earth Science and Technology (SOEST) University of Hawaii 1680 East-West Rd Honolulu, HI USA [email protected] Gerald G. Garland Department of Geography and Urban Planning University of United Arab Emirates P.O. Box 17771 Al Ain United Arab Emirates [email protected] Ioannis Georgiou Department of Earth and Environmental Sciences Pontchartrain Institute for Environmental Sciences University of New Orleans New Orleans, LA 70148 USA [email protected] P. Giresse Laboratoire d’Étude des Géo-Environnements Marins Université de Perpignan 52, Av. Paul Alduy 66860 Perpignan France A. Goudie School of Geography and the Environment Oxford University South Parks Road, Oxford, OX1 3QY United Kingdom [email protected]

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Olavi Granö Department of Geography University of Turku Turku 20014 Finland [email protected]

David Hopley School of Earth and Environmental Sciences James Cook University Townsville, Queensland 4811 Australia [email protected]

André Guilcher†

John R. C. Hsu Department of Marine Environment and Engineering National Sun Yat-sen University 70 Lienhai Road, Gushan District Kaohsiung City 80424 Taiwan [email protected]

Mario Gutiérrez-Estrada Institute of Marine Sciences Mazatlan Mexico P.T. Hanamgond Dept. of Geology GSS College Tilakwadi, Belgaum India Jim Hansom Department of Geographical and Earth Sciences University of Glasgow Glasgow, G12 8QQ Scotland UK [email protected] Miles Hayes Department of Geology University of South Carolina Columbia, SC 29208 USA [email protected] Terry Healy Department of Earth and Ocean Sciences University of Waikato Hamilton 3240 New Zealand [email protected] Irénée Heyse Department of Geography University of Ghent Krijgslaan 281-S8 Ghent Belgium [email protected]

Federico Isla Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Centro de Geología de Costas y del Cuaternario Funes 3350, Mar del Plata Argentina [email protected] W. Jara† Chen Jiyu State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China [email protected] Pavel Kaplin Department of Geology Odessa State University Odessa 270000 Ukraine [email protected] Tormod Klemsdal Department of Geosciences University of Oslo P.O. Box 1047 Blindern 316 Oslo Norway [email protected]

List of Contributors

Kazuyuki Koike Department of Geography Komazawa University Komazawa, Setagaya-ku Tokyo 154 Japan Paul D. Komar College of Oceanic & Atmospheric Sciences Oregon State University Corvallis, OR 97331 USA [email protected] Are Kont Institute of Ecology Tallinn University Narva mnt 25 Tallinn Estonia [email protected] Mark Kulp Department of Earth and Environmental Sciences University of New Orleans 2000 Lakeshore New Orleans, LA USA [email protected] M. Larid University Center Stambouli Mustapha of Mascara Route de Mamounia 29000 Mascara Algeria Eggert Larusson School of Education University of Iceland Stakkahlíð IS 101 Reykjavík Iceland [email protected] Jean-Michel Lebigre Université de la Nouvelle-Calédonie BPR4 98851 Nouméa Cedex New Caledonia Canada [email protected]

Tommy Lindell Centre for Image Analysis Swedish University of Agricultural Sciences Uppsala University Box 337 Uppsala Sweden [email protected] Svetlana Lukyanova Faculty of Geography M.V. Lomonosov Moscow State University Moscow Russia [email protected] Kyaw Saw Lynn Anglo-Chinese Junior College 25 Dover Close East Singapore [email protected] Molly McGraw Southeastern Louisiana University Hammond Louisiana, LA 70402 USA [email protected] [email protected] Michael Miner Department of Earth and Environmental Sciences University of New Orleans New Orleans, LA 70148 USA Guy De Moor Geological Institute University of Ghent Krijgslaan 281-S8 Ghent Belgium Carlos Morais Laboratório Nacional de Engenharia Civil Av do Brasil, 101 Lisboa Portugal

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Maria Eugénia Soares de Albergaria Moreira Quinta das Mós - Rua da Cruz nr 5 Zibreira da Fé Sobral de Monte Agraço Portugal 2590–293 [email protected]

Patrick Nunn Department of Geography University of the South Pacific P.O. Box 1168 Suva, Fiji Islands [email protected]

Jack Morelock Department of Marine Sciences University of Puerto Rico at Mayagüez P.O. Box 908 Lajas, PR Puerto Rico

Francis Ojany Department of Geography University of Nairobi 30197 Nairobi Kenya 00100-GPO [email protected]

Robert Morton United States Geological Survey 10100 Burnet Rd., Bldg. 130 Austin, Texas 78758 USA [email protected] Dieter Muehe Department of Geography University of Copenhagen Øster Voldgade 10 DK-1350 Copenhagen Denmark [email protected] G. N. Nayak Department of Marine Science Goa University Taleigao Plateau Goa 403206 India [email protected] [email protected] Yaacov Nir 15 Shimeoni st. Rehovot 76248 Israel [email protected] John Norrman Department of Earth Sciences Uppsala University Villav. 16 Uppsala Sweden

Otto Ongkosongo Lembaga Oseanologi Nasional Jl. Pasir Putih I Ancol Timur Jakarta Utara Indonesia [email protected] Mario Arturo Ortiz-Perez Institute of Marine Sciences Mazatlan Mexico Kaarel Orviku GONSIORI str. 24-23 Talinn Estonia [email protected] Ervin Otvos Department of Coastal Sciences Ocean Springers Campus University of Southern Mississippi Ocean Springs MS 39566 USA [email protected] André Ozer Département de Géographie Physique Université de Liège 2 Allée du 6 Août Liège Belgium [email protected]

List of Contributors

Wong Poh Poh Department of Geography National University of Singapore 21 Lower Kent Ridge Road Singapore 119077 [email protected]

Vidal Taype Ramos Geophysics Commission Pan American Institute of Geography and History Av. A. Aramburú 1190-1198 Lima 34 Peru

Wayne Pollard Department of Geography McGill University 805 Sherbrooke St. W. Montreal, Quebec Canada H3A 2K6 [email protected]

Stanley R. Riggs Department of Geology Thomas Harriot College of Arts and Sciences East Carolina University 1002 Bate Building East Fifth Street Greenville, NC USA [email protected]

Alexey Porotov Faculty of Geography Lomonosov Moscow State University GSP-1, Leninskie Gory Moscow, 119991 Russian Federation [email protected] Jorge Pousa Consejo Nacional de Investigaciones Científicas y Técnicas (CONICET) Universidad Nacional de La Plata (UNLP) Calle 64 (entre 119 y 120) La Plata, Buenos Aires Province Argentina [email protected]

William Ritchie AICSM University of Aberdeen Elphinstone Road Aberdeen AB24 3UF Scotland UK [email protected] Ridel Rodríguez Apartado postal 127 Holguín, C.P. Cuba 80 100 [email protected] [email protected]

Norbert Psuty Institute of Marine and Coastal Sciences Rutgers University 74 Magruder Road Sandy Hook, NJ USA [email protected]

Joanna Rotnicka Institute of Geology Adam Mickiewicz University Maków Polnych 16 Poznañ Poland [email protected]

Garry Ramedine Université des antilles et de la guyane Campus de Fouillole, BP 250 Pointe-à-Pitre, Grande-Terre Guadeloupe

Karol Rotnicki Institute of Geoecology and Geoinformation Adam Mickiewicz University Ul. Dziegielowa 27 Poznan Poland [email protected]

Wilson Ramirez Geology Department University of Puerto Rico at Mayagüez P.O. Box 9017 Mayagüez, PR Puerto Rico

Teh Tiong Sa 19 Jalan Jintan #12-25 Singapore [email protected]

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Pascal Saffache Université des antilles et de la guyane Campus de Fouillole, BP 250 Pointe-à-Pitre, Grande-Terre Guadeloupe [email protected] George Said Favorita Sir Paul Boffa street Victoria Gozo, Malta [email protected] Paul Sanlaville Le Sanderey St. Michel de Villadeix F-24380 France Anja Scheffers School of Environmental Science and Management Southern Cross University P.O. Box 157 Lismore, NSW 2480 Australia [email protected] John Schembri Favorita Sir Paul Boffa street Victoria Gozo, Malta Enrique Schnack Laboratorio de Oceanografia Costera Facultad de Ciencias Naturales y Museo Universidad Nacional de La Plata C.C. 45, 1900 La Plata, Buenos Aires Argentina [email protected] Maurice Schwartz Department of Geology Western Washington University Bellingham, Washington 98225 USA [email protected]

Klaus Schwarzer Coastal and Continental Shelf Research Institute of Geosciences University of Kiel Olshausenstrasse 40 Kiel Germany [email protected] Andrei Selivanov† Yurii Shuisky Physical Geography Department Mechnikov’s National University of Odessa 2 Dvoryanskaya St. Odessa-82 Ukraine [email protected] Scott Smithers School of Earth and Environmental Sciences Faculty of Science and Engineering James Cook University Queensland 4810 Australia Rodman E. Snead Department of Geography University of New Mexico Albuquerque New Mexico USA Horst Sterr Institute of Geography University of Kiel Olshausenstrasse 40 Kiel Germany [email protected] Thomas Terich Center for Geography and Environmental Social Sciences Huxley College of Environmental Studies Western Washington University Bellingham, WA 98225 USA [email protected]

List of Contributors

Bruce Thom School of Geosciences The University of Sydney Australia [email protected] Hannes Tõnisson Institute of Ecology Tallinn University Tallinn Estonia [email protected] J. Turenne Centre d’Ile-de-France 32, avenue Henri Varagnat 93143 Bondy cedex France C. R. Twidale Department of Geography University of Adelaide Adelaide, South Australia 5005 Australia [email protected] Etop Usoro Department of Geography and Regional Planning University of Uyo 1, Ikpa Road, P.M.B. 1017 Uyo, Akwa Ibom State Nigera Don Vermeer Louisiana State University Baton Rouge, LA 70803 USA

H. Viles School of Geography and the Environment Oxford University South Parks Road, Oxford, OX1 3QY United Kingdom [email protected] Jesse Walker Department of Geography Louisiana State University Baton Rouge, Louisiana 70803 USA [email protected] Gabriel Yong Department of Geography Universiti Brunei Darussalam Jalan Tungku Link Gadong Negara Brunei Darussalam BE1310 [email protected] Dai Zhijun State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China Yu Zhiying State Key Lab. of Estuarine and Coastal Research East China Normal University Northern Zhongshan Road 3663 Shanghai 200062 P.R. China

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1.0 USA – Editorial Introduction

The major geomorphological divisions of the United States, the Rocky Mountain ranges, the Great Plains and the Appa­lachian ranges, do not have a major influence on coastal features. The Pacific coast follows the alignment of the Coast Ranges on the western side of the Rocky Mountains, the Great Plains run southward to the lowlying coast of the Gulf of Mexico, and the Appalachians extend north-eastward to the coast of Maine. Useful background is provided by Shepard and Wanless (1971), Sherman (2005), Thornbury (1965) and Graf (1987).

1.  Alaska and Pacific Coast The Alaskan coast (> Alaska) is long, varied, and irregular, some of it steep and mountainous. It has developed where the northern extension of the Rock Mountain ranges swings westward, and is separated by valleys along intermontane basins. It has been strongly glaciated, and some of the fiords are still occupied by glaciers, as in Nunatak Fiord, though they have been receding. Tectonic activity has been vigorous along the Alaskan Peninsula and the Aleutian Island, where there are several active volcanoes along the plate convergence. South-eastern Alaska is rising rapidly, locally up to 3.5 cm/year. Mean spring tide ranges are small on the Alaskan Peninsula and the Aleutian Islands, but increase in the Gulf of Alaksa to more than 10 m near Anchorage. Much of the Pacific coast ( > Washington,  > Oregon, > California) is steep and high, and subject to continuing tectonic activity, including movements along the San Andreas Fault and tectonic disruptions associated with the subduction zone offshore, where the Pacific Plate is passing beneath the North American Plate. The continental shelf is narrow ­(20–30 km), and the coast is exposed to strong wave energy from ocean swell transmitted across the Pacific as well as storms in coastal waters. Mean spring tide ranges are generally small ­(1–3 m), but the coast is subject to occasional tsunamis, including those generated by tectonic disruptions along the offshore zone of converging plates. The steep coasts are generally vegetated slopes descending to rocky cliffs. They are interrupted by the mouths of numerous valleys, with marine inlets or coastal lagoons

(bays) enclosed by spits and barriers, mainly sandy. The spits and barriers are generally capped by dunes, and dunes become extensive in some sectors, ­notably in Central Oregon, where they have advanced several kilometres inland. Fluvial sediment from the Columbia River has been the major source of sandy beach material, which has drifted from the river mouth both north into Washington and south into Oregon. Tectonic activity has resulted in the uplift of marine terraces. These are prominent in California, where they are found at various levels up to 411 m on the Palos Verde Peninsula, rising to 600 m near Santa Barbara, then falling to 200 m near Big Sur. San Francisco Bay is a large basin dating from Pliocene times that has been invaded by the sea. Coastal rock formations from Washington State south to northern California include mélanges of sedimentary, metamorphic, and volcanic rocks assembled in a tectonically disturbed environment. The numerous stacks and islands off this coast, sometimes extending several kilometres seaward, are largely the more resistant residuals of these formations, isolated in the course of long-term coastline recession. The recession of steep, forested coastal bluffs in Washington and Oregon has been intermittent, the cumulative result of recurrent landslides and episodes of marine erosion during tsunamis. On the Olympic Peninsula in Washington, the coast had been influenced by glacial and periglacial processes during cold phases of the Pleistocene. One indication is the presence of gravelly beaches.

2.  Gulf of Mexico Coast Several rivers have fed fluvial sand and silt to the Gulf of Mexico coast, the largest being the Mississippi, which has built a major composite delta. In > Texas, there are large sandy dune-capped barriers, backed by beach ridges, lagoons, and marshes and deltaic sectors such as that at the mouth of Brazos River. The coastal plain is wide in > Louisiana, fringed seaward by beaches and cheniers perched on alluvial flats. The tide range on the Gulf coast is small, and the coast has been shaped by generally gentle Caribbean waves and occasional hurricane surges. East of the Mississippi delta, wave energy is relatively low, but

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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1.0    

USA – Editorial Introduction

s­ ufficient to have formed beaches and barrier islands on the coasts of > Mississippi and > Alabama. Further east, it diminishes along the Florida coast (> Florida), and is exceptionally low in Apalachee Bay. The western coast of the Florida peninsula is low-lying with extensive marshes and mangroves threaded by inlets.

3.  Atlantic Coast Much of the Atlantic coast of the United States is a lowlying trailing edge coast with a continental shelf up to 100 km wide. Long sandy barriers and dune-capped barrier islands extend along the Atlantic coast from Florida to New Jersey (> Florida, > Georgia, > South Carolina, > North Carolina, > Central Atlantic Coast, > New York and New England, > Great Lakes (USA) and > Hawaii), backed by intermittent lagoons. The tide range is small (1–2 m) and wave action mainly south-easterly from Atlantic storms; hurricanes have occasionally produced surges causing erosion and flooding. In the absence of rivers yielding sediment to the sea, the beaches, barriers, and dunes are predominantly quartzose, with local accessions of shelly material. Longshore drift is predominantly northward, forming compound recurved spits such as Cape Henry and Sandy Hook, but there are episodes of southward drifting, the alternations contributing to the shaping of the major cuspate forelands, Cape Canaveral, Cape Fear, Cape Lookout, and Cape Hatteras. Each of these has beach and dune ridges marking stages in evolution. There are large estuarine embayments such as Delaware Bay, Chesapeake Bay, and the barrier-enclosed Pamlico Sound. Many of the coastal lagoons and inlets have muddy areas, with salt marshes. In Florida, there are intertidal sabellarid reefs, and to the south coral reefs. Beaches are more calcareous, with sediment derived from these, and mangroves appear in the salt marshes. In New England (> New York and New England) to the north the coast is hilly and irregular on structures associated with the northern Appalachians. Glaciation extended as far south as New York, and has left numerous valley-mouth inlets, submerged during the Holocene marine transgression, many of them now blocked by barriers or sand and shingle to form lagoons known as ponds. There are also glacial outwash features and drumlins, as in Boston Harbour, and moraines, including Long Island, New York. Reworked glacial drift has produced sand, pebbles, cobbles, and some boulders. There are sand and shingle spits and beaches and sand dunes, as on Long Island. Staten Island, New York, is a morainic feature close to the limit of glaciation.

4.  Great Lakes The Great Lakes (> Great Lakes (USA)) are a group of linked freshwater lakes discharging into the St. Lawrence River. Although they form a natural unit, it is necessary to split them because the boundary between the United States and > Canada passes through Lakes Superior, Huron, Erie, and Ontario (so that their southern coastlines are in the United States), whereas Lake Michigan is entirely within the United States. The present Great Lakes began to form about 15,000 years ago when the Wisconsin ice sheet was melting, and water accumulated between the receding ice margin and a drainage divide to the south. The shaping of the lakes was influenced by the topography exposed as the ice retreated and by uplift as the ice load decreased (postglacial isostatic recovery). Lake Erie is 174 m above mean sea level, and outflow at Buffalo to the Niagara River descends Niagara Falls to Lake Ontario, 75 m above sea level, and then along St. Lawrence River to the sea. The lakes are almost tideless, but show fluctuations in level related to atmospheric pressure, wind stress, precipitation and evaporation, and freshwater inflow and outflow regimes. Seasonal fluctuations are of the order of 0.35 m, but longer-term oscillations of up to 2 m have occurred. Waves are generated by local winds, and may be up to 4 m high during storms. Parts of the lakes coastline are rocky on Pre-Cambrian and Palaeozoic igneous and sedimentary formations, but glacial drift deposits are extensive, and in places these have been cut back as cliffs. Sand and gravel derived from glacial drift form narrow ( Hawaii) is of volcanic origin, formed by extrusion of basalt lava along a midPacific zone where a hot spot has migrated NW-SE beneath the Pacific plate. Successively formed volcanic structures have been dissected by Pacific Ocean swell and storm waves to form cliffs, some of which are high and ­spectacular. Coastal landforms include craters invaded by the sea. Boulders, gravel, and sand derived from the volcanic rocks

USA – Editorial Introduction  

form beaches. Wave energy is high, concentrated on these microtidal coasts, and wave break heavily on surf beaches. Coral reefs fringe some sectors of coast, and coralline sand and gravel is a constituent of many beaches, along with shelly and algal sediment. Some calcareous beaches are backed by dunes and dune calcarenite. There are emerged reefs and associated shore features resulting from tectonic uplift and sea level variations, and the islands have been subject to recurrent tsunamis, originating from earthquakes and volcanic eruptions around the rim of the Pacific basin. A major tsunami in 1946, caused by an earthquake off the Aleutian Islands, produced giant waves on Hawaiian coasts, as did the tsunami resulting from the Chilean earthquake in 1960.

   1.0

References Graf WL (ed) (1987) Geomorphic systems of North America. Geological Society of America, Boulder, CO Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw Hill, New York Sherman DJ (2005) North America coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 721–727 Thornbury WD (1965) Regional geomorphology of the United States. Wiley, New York

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1.1 Alaska Jesse Walker · Molly McGraw 

1.  Introduction Alaska, which spans 20° of latitude and 50° of longitude, has a coastline about 10,686 km in length. Because of its position and shape, it faces three distinct bodies of water, the Arctic Ocean on the north and northwest, the Pacific Ocean on the south, and the smaller, semi-enclosed Bering Sea on the west. Alaska occupies the northwest extension of four of North America’s major physiographic divisions (­Wah­rhaftig 1965): the Interior Plains, the rocky Mountain System, the Intermontane Plateaus, and the Pacific ­Moun­tain System

(> Fig. 1.1.1). All of these physiographic divisions extend across the Alaskan coast and each has an important bearing on coastal geomorphology and geology (> Fig. 1.1.2) and coastline character (> Fig. 1.1.3). The Interior Plains in Alaska, known as the Arctic Coastal Plain, is composed almost entirely of Quaternary sediments most of which are unconsolidated and some aeolian. The coastal zone of the Brooks Range (the Alaskan portion of the Rocky Mountain System) consists of Mesozoic and Palaeozoic sedimentary and metamorphic rocks. Most of the Alaskan Bering Sea Coast belongs to the Intermontane Plateaus. It is composed mainly of

⊡⊡ Fig. 1.1.1 Physiographic divisions of Alaska.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Alaska

⊡⊡ Fig. 1.1.2 Coastal geology of Alaska.

Quaternary sediment although there are several sections where Mesozoic and Palaeozoic sedimentary and metamorphic rocks predominate. In addition, a few locations have Quaternary and Tertiary volcanics and intrusive igneous rocks of Tertiary age. The geology of the Pacific Mountain System is most diverse. Virtually, all rock types are present. In the Aleutian Islands, Quaternary volcanics predominate; along the central Gulf of Alaska coast, sedimentary and metamorphic rocks of Quaternary, Tertiary, and Mesozoic ages are all present; and along the lengthy coastline of southeast Alaska there is a mixture of Tertiary and Mesozoic intrusives and Mesozoic and Palaeozoic sedimentary and metamorphic rocks. The continental shelf of Alaska varies greatly in width being quite narrow along the Pacific Ocean and quite wide in the Bering Sea and the Arctic Ocean. With a width of only about 150 km opposite Demarcation Point (the

northwest corner of Alaska), it widens westward toward the Siberian Coast. The northern half of the Bering Sea is sufficiently shallow so that during glacial epochs, it was subaerial, connecting eastern Siberia with western Alaska. Along the Pacific coast of Alaska, the continental shelf is bordered seaward by the Aleutian Trench, which results from the subduction of the Pacific Plate beneath the North American Plate. These contrasts are also evidenced in the amount of tectonic activity affecting the coastal zone.Along the Arctic Ocean and Bering Sea coastlines, there is minimal earthquake and volcanic activity, whereas along the Gulf of Alaska coast, it tends to be intense. The climate of coastal Alaska varies with latitude. Three distinct types of climate, polar, continental, and marine, are represented along the Alaskan coastline. Polar climatic conditions are characterised by average low ­temperatures less than −10°C and occur along the Arctic

Alaska

1.1

⊡⊡ Fig. 1.1.3 Coastline character of Alaska.

and Chukchi Seas, roughly from Demarcation Point to Point Hope. The Arctic Ocean has minor influence on temperature moderation and precipitation along the coast. The prevailing winds are easterly and are strong along the coast, but decrease inland. Temperatures within this region range from a mean annual high of −2.3°C in Kotzebue to a mean annual low of −15°C along the Arctic Ocean coast. Total average annual snowfall ranges from 760 mm along the Arctic coastline to 1,260 mm in Kotzebue. The continental climate is characterised as wet with cold winters having average temperatures of less than −3°C. Generally, both summer and winter temperatures are extreme and precipitation is light. Except for a few isolated locations, surface winds are light and from the north and east. These conditions occur along the southern Chukchi Sea and Bering Sea, from Point Hope to Unimak Island. Temperatures within this region range from a mean annual

high of 5.6°C at Port Heiden to a mean annual low of −9°C in Wales. Average annual snowfall ranges from 970 mm in Wales to 1,670 mm at Moses Point. The marine climate is characterised by wet, mild winters and cool summers. This type of climate is found along the southern coast, from the Aleutian Islands to the Alexander Archipelago. Temperatures vary little from place to place in southeast Alaska, but precipitation changes based on exposure to the influence of the warm Japan Current and by the proximity of the Gulf of Alaska and by altitude. Strong northerly winds occur along the Gulf coast. Winds are generally from the southeast in the Alexander Archipelago. Most of the annual precipitation is in the form of rainfall, which is heavy for most of the region. Average annual snowfall ranges from 100 mm in Sitka to 6,360 mm at Whittier. Temperatures are moderate with an average annual range of −2° to 6°C in Anchorage and from −4° to 10°C in Sitka.

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Alaska

Tides vary significantly along the Alaskan coast. The coastal tide range generally increases from north to south. The tide range for Barrow, on the Arctic Ocean, is about 0.4 m; that at Port Heiden, on the Bering Sea coast, is about 12 m. High tide ranges are found along the Gulf of Alaska coast and in the Alexander Archipelago. The maximum diurnal tide range for Anchorage is about 9 m and at Whittier it is about 12 m. Sea ice affects most of coastal Alaska. The Arctic Ocean remains frozen roughly from October to June. The ice may reach thicknesses of 200 cm, and 2–3 m pressure ridges are common. The Bering Sea freeze-up generally begins in mid-November. Ice break-up occurs in late May and early June. Ice thickness varies with latitude along the coast. Maximum thicknesses of 161 cm and 82 cm have been calculated for the Kotzebue and Kuskokwim areas, respectively. Bering Sea ice typically extends to the upper portion of Bristol Bay. Sea ice is generally confined to the Cook Inlet along the Gulf of Alaska coast and begins forming in November and is completely broken up by mid-April. Two of the major conditions affecting coastline development in Alaska are permafrost and glaciers. Permafrost is especially significant along the northern coast of Alaska while glaciers are common along the southern coast. Although only the southern part of Alaska and the Aleutian Islands are free of permafrost, the continuous variety (the type that impacts coastal environments) is confined to the coastline north and east of the Bering Sea. In that area permafrost, with its included ice wedges, affects the character and amount of coastal erosion. In the zone of discontinuous permafrost, coastline erosion frequently exposes relict permafrost patches. Along the north coast of Alaska, permafrost that formed during lower sea stands is present in nearshore, sub-sea locations. Except for a few glaciers present in the Brooks Range of northern Alaska, glaciers today are confined to southern Alaska where many still impact the coastline. Most of the Alaska coast, even during the period of maximum glacial extent, was untouched directly by glaciers. Along the Arctic Ocean and Bering Sea coasts, sediment from glaciers (as for example via the Colville River draining the Brooks Range glaciers) reached the sea. During the Pleistocene, glaciers along the Gulf of Alaska coast advanced out over the continental shelf. Today, they are confined to the coastal plain or coastal mountains. Many glaciers in southeastern Alaska have their calving termini in fiords (Molnia 1986; Trabant et al. 2003). In Alaska, geologic and oceanic settings combine to provide five coastal units: Arctic, Bering Sea, Aleutian Islands, Gulf of Alaska, and Southeast. Within these five zones, coastal processes and resultant landforms are

numerous and highly varied. Some forms are small and ephemeral such as sea-ice kettles that develop seasonally along the Arctic coast; some are large and active like the Yukon River delta; some, such as the lava coasts in the Aleutian Islands chain, are the result of catastrophic events; others, like the extensive tidal flats in Cook Inlet, are subjected to daily modification and yet others are large and durable as exampled by Chatham Strait, a 400 km long fiord in southeast Alaska.

2.  The Arctic Coast The Arctic coast of Alaska extends from Demarcation Point on the Canada–Alaska border west to Point Barrow and then southwest to Cape Prince of Wales. Point Barrow, the location of the most northern town in the United States, is also the dividing point between the Beaufort Sea to the east and the Chukchi Sea to the west. From the international border west to Point Barrow, the coast is characterised by low, ice-rich cliffs, except where interrupted by the numerous rivers that flow across the coastal plain. Some of these cliffs face the open sea although most are separated from it by a series of shallow lagoons and low barrier islands. Most cliffs are 1–5 m high although toward the east some are as much as 9 m high (Walker 1991). These cliffs are bordered by narrow beaches. The barrier islands, most numerous along the eastern half of the Beaufort Sea coast, are generally narrow and short and separated by wide inlets. The western half of the Beaufort Sea coast is irregular and is composed of a number of large bays. One large delta, the Colville River delta, and a number of small deltas are present along the coast (Walker 1974). Coastal erosion yields sediment and carbon inputs (Jor­genson and Brown 2005). The Chukchi Sea coast generally has higher relief than the Beaufort Sea coast. Southwest of Point Barrow, tundra cliffs are usually between 10 and 18 m high (Harper 1978) although at Cape Lisburne, where the Brooks Range extends to the shore, they are more than 300 m high (> Fig. 1.1.4). An unvegetated talus slope extends down into the sea. Lengthy spits and barrier islands, separated by narrow inlets, occur along the middle section of the coast between Point Barrow and Cape Lisburne and along the northwest facing portion of Seward Peninsula. Shorter barrier bars are present along much of the rest of this coast. In some places, as at Cape Krusenstern and Point Hope, extensive beachridge systems have developed (> Fig. 1.1.5). Point Hope is a large cuspate formation that has grown out into the Chukchi Sea, flanked by long barriers that enclose lagoons and a swampy coastal plain with rounded lakes. One of

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⊡⊡ Fig. 1.1.4 Cliffs south of Cape Lisburne. (Photo by J.L. LaBelle, Courtesy ARLIS.)

⊡⊡ Fig. 1.1.5 Beach ridges at Point Hope. (Photo by C.D. Evans, Courtesy ARLIS.)

the most distinctive features of the Arctic Coast is a large embayment located north of Seward Peninsula. Sediment transported by the Noatak and Kobuk, two of the largest rivers in Alaska, is rapidly filling this embayment. In arctic Alaska, a low tide range and a long-lasting seaice cover combine to limit the amount of ­wave-generated

coastal erosion. Even during summer, especially along the Beaufort Sea coast, sea-ice is seldom far from shore and effectively reduces fetch and dampens wave action. Nonetheless, coastal retreat can be rapid. When the sea-ice retreats from the shore, usually during summer and fall, storms may cause severe erosion. The two towns of Barrow

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and Wainwright are exposed to a long fetch when the sea-ice retreats from the shore. Nearly all of the arctic coast is within the zone of continuous permafrost and exhibits such permafrost-related forms as ice wedges, ice-wedge polygons, and thaw lakes and permafrost-related processes as thermal melting and thermal erosion. Ice wedges, which may occupy as much as 25% of the upper several metres of the ground, and icewedge polygons, which form a dense network over much of the Arctic Coastal Plain (> Figs. 1.1.6 and >1.1.7), are frequently subjected to wave attack. Differential thaw and erosion result in the undercut, low-cliff, serrated-type coastline that characterises much of the Arctic. Thaw lakes when tapped by ocean waves become part of the shallow foreshore and provide a coastline that is usually arcuate in form. The Cape Lisburne area is dominated by cliffs as much as 300 m high and represents the western end of the Brooks Range. About 50 km south of Cape Lisburne is Point Hope, a 15 km long sand spit. An extensive series of beach ridges has built up at Cape Krusenstern south of Point Hope.

3.  The Bering Sea Coast The Bering Sea Coast extends south from Cape Prince of Wales to the west end of the Alaska Peninsula. It is

dominated by two large embayments (Norton Sound to the north and Bristol Bay to the south) that are separated by the extensive Yukon–Kuskokwim deltaic complex and the southwest extension of the Kuskokwim Mountain Range. Norton Sound, with a relatively smooth coastline, has narrow beaches that are backed by glaciated hills. The Yukon–Kuskokwim delta, a large triangular area with an 800 km long coastline, consists of many lakes, marshes, and abandoned distributaries. Presently, the Yukon River, which contributes the majority of the sediment that reaches the Bering Sea, discharges north into Norton Sound where interdistributary mudflats are extensive. South of the Yukon, the coast is dominated by beach ridges, many of which are undergoing rapid erosion. The Kuskokwim River, in contrast to the Yukonl, flows into a 100 km long estuary. The Yukon–Kuskokwim delta is the seventh largest delta on earth, covering about 54,000 sq km. It is a vast plain with both wet and dry tundra and numerous lakes. Tidal flats, typically 100–1,000 m wide, flank the Kuskokwim delta (Dupre 1977). Although most of the coast between the Yukon and Kuskokwim is made of deltaic deposits, that in the SW part of the deltaic complex, where Quaternary volcanic activity formed hills, includes gravel beaches and rocky headlands. Nelson Island near the Kuskokwim Delta and Nunivak Island about 50 km offshore, have such coastlines.

⊡⊡ Fig. 1.1.6 Ice-wedge polygons during snow melt and break-up.

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⊡⊡ Fig. 1.1.7 Progressive ice-wedge degradation in the Colville River delta.

In contrast, much of the coast of Saint Lawrence Island has low sandy beaches often backed by sand dunes. On its south coast sand and gravel barriers enclose Koozata Lagoon and Sekinak Lagoon, their inner shores having numerous cuspate spits marking early stages in lagoon segmentation. South of the Kuskokwim Estuary, the Kuskokwim Mountains meet the sea with high rocky cliffs and numerous stacks. Most of the Bristol Bay coast is composed of glacial and fluvial sediment and has sandy beaches, extensive mudflats, and lagoons. The tide range along the Bering Sea coast, although variable, in general increases from less than 1 m in Norton Sound to more than 5 m in Bristol Bay. In contrast, the amount of sea-ice decreases southward; ranging from a continuous cover during more than half of the year in the north to only occasional occurrences, even in mid-winter, in the south. Thus, whereas the coastline north of the Kuskokwim is protected by sea-ice during winter, that in Bristol Bay is subjected throughout the year to many of the storms that cross from the Pacific Ocean into the Bering Sea. As a result, most of the unconsolidated coasts in Bristol Bay have been eroded into low cliffs.

4.  The Aleutian Islands The Aleutian Islands begin at the west end of the Alaska Peninsula and extend in a gentle arc for nearly 2,000 km

across the north Pacific Ocean to Attu Island. The islands are the subaerial portions of a 10–100 km wide ridge that flanks the western half of the Aleutian Trench. This trench (over 3,000 km long and up to 7,500 m deep) is the subducting edge of the Pacific Plate. Active subduction is responsible for the occurrence of frequent earthquakes in south Alaska and for the creation and maintenance of the 60 centres of volcanic activity that extend from near Anchorage to the west end of the Aleutian Islands. Of these 60 centres, more than 40 have been active since 1760 (Selkregg 1976). Many of the coasts in the Aleutian Islands are composed of lava that flowed into the sea or of ash that has fallen on the shore or into the water and then washed on shore; other coasts have been eliminated or altered by explosive eruptions. One of the best examples of such a coast is that on Bogoslof Island, which lies near Bering Canyon (the longest submarine canyon in the world). The volcanic history of Bogoslof Island has been followed since 1769 (Shepard and Wanless 1971). The coastlines of most of the islands are irregular. Rocky cliffs alternate with boulder beaches although in some protected coves sandy beaches are present. Abrasion platforms are found at present-day sea level as well as at various heights below and above (up to 200 m) sea level. In the Aleutian Islands, wind-generated waves, which are high throughout the year, combine with high tide ranges to cause extensive coastal erosion. In places, erosion rates are sufficient to have produced stacks along the

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shore and hanging valleys at the mouths of many streams. Other factors affecting coastal development in the islands are the vast quantities of kelp, which grows well in the turbulent water (and diminishes wave energy) and the occasional tsunami some of which have carried driftwood to heights of more than 30 m (Black 1974).

5.  The Gulf of Alaska Coast The Gulf of Alaska Coast, which extends in a clockwise direction from the eastern Aleutian Islands to Cape Spencer, includes a number of distinctive coastal types. The Pacific facing portion of the Alaskan Peninsula, in contrast to its low-lying Bering Sea counterpart, is highly irregular and has numerous rocky headlands that alternate with formerly glaciated inlets. Many of the headlands are represented offshore by rugged islands and rocky stacks. One of the largest islands in Alaska is Kodiak, which is separated from the northern part of the Alaska Peninsula by a 50-km wide strait. The coastline of Kodiak is also irregular and rugged. Fiords, especially along its northwest coast, are numerous. Cook Inlet, a deep structural basin over 300 km long, is a tidal estuary that opens into the Gulf of Alaska and has some of the world’s highest tide ranges. Near its head, extensive tidal flats and marshlands occupy the coastal edge of a low coastal plain. The plain narrows gradually on the northern side of Cook Inlet and toward its mouth merges into mountains that descend directly into the sea. In contrast, on the south side of Cook Inlet, pocket beaches alternate with low cliffs along most of its extent. Associated with the high tide range of Cook Inlet are strong tidal currents; the two help keep ice from forming a solid cover over the Inlet during winter. The coast from the mouth of Cook Inlet east around Prince William Sound is deeply indented with fiords and possesses many ice-scoured headlands and islands. Much sediment is transported into the fiords by the short, turbulent streams that originate in the alpine glaciers that are still present in the coastal mountains. From Prince William Sound southeast to Cape Spencer, the coastline is smooth in outline and is characterised by deltas, beach and dune ridges, outwash plains, moraines, and glaciers. In some locations, tidewater glaciers are present and represent some of the most rapidly changing coastlines to be found in Alaska. One of the major features of this coast is Malaspina Glacier, said to be the world’s largest piedmont glacier (Hayes et al. 1976). Although all of the Pacific Ocean coast of Alaska may experience earthquakes, some of the most severe occur around the Gulf of Alaska and many have a direct effect on

the coastline. The March 27, 1964 Alaskan Earthquake, the most intense ever recorded in North America, depressed, elevated, or otherwise modified some 16,000 km of coastline (Shepard and Wanless 1971). Montague Island was raised by nearly 10 m; other coastlines were lowered, some by as much as 3 m. Homer Spit, in Kachemak Bay, sank nearly 2 m, and spruce forests on low-lying coasts in Turnagain Inlet were replaced by salt marshes. There was slumping on coastal bluffs near Anchorage, forming debris cones that extended up to 600 m into the sea. Such sudden emergence establishes a new zone on which subaerial processes can operate, resulting in a replacement of marine biota, and shifts seaward the zone of marine processes. Submergence, on the other hand, leads to the incursion of saltwater into coastal lakes and marshes, the loss of beaches, spits, and bars, and the subjection of former subaerial surfaces to wave attack. Much of the Gulf of Alaska coast shows evidence of these changes (Waller 1966; Hamilton and Shennan 2005). Earthquakes are a major (although not the only) cause of landslides in the coastal zones of Alaska. Some landslides occur offshore especially at the front of the subaqueous portion of deltas, whereas others happen along the steep subaerial slopes leading down to the shore. The most famous slide was triggered by a July 9, 1958 earthquake near Lituya Bay. A horizontal movement of 6.6 m and vertical movement of 10.7 m caused a rockslide that dumped more than 30 million m3 of rock into the bay. The wave that the rockfall created swept 530 m up the opposite side of the bay destroying all vegetation in the process. Trim­ lines from such waves are conspicuous along its coast. Earthquakes also trigger calving of ice at the terminus of glaciers; for example, during the 1958 earthquake Turner Glacier at the head of the Yakutat Bay was reduced in length by 750 m. Hubbard Glacier is another example of calving glaciers along this shoreline (>Fig. 1.1.8).

6.  The Southeast Alaska Coast The Southeast Alaska coastline extends from Cape Spencer south to Tongass at Dixon Entrance on the Alaska–Canada border. Although the general length of the Southeast Alaska Coast, as given in the United States Coast Pilot (1981), is only 450 km (4.3% of Alaska’s total), the tidal shoreline, is more than 45 times longer. Such a variation is because of the large number of inlets, straits, bays, fiords, and islands that characterise the region. The coast of most of southeast Alaska is rugged with steep cliffs many of which descend far below sea level. However, there are low coastal plains near the head of some of the fiords and

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⊡⊡ Fig. 1.1.8 Hubbard Glacier. (Photo by W.D. Harrisson, Courtesy ARLIS.)

strandflats and reefs border many islands and headlands, especially those exposed directly to the ocean. Many of the geomorphic details are closely related to the intensity of wave and tidal action and lithology. Most of the coastline is protected from the open sea and thus not generally subjected to the high energy storm waves that characterise this part of Alaska. Nonetheless, tides of 7 m and more in some of the fiords and bays are responsible for strong tidal currents and intensive scouring action. In contrast, the gross geomorphic character of southeast Alaska is the result of the glacial modification of a tectonically active, complex mountain terrain. At the maximum of the Plaeistocene Glaciation, ice covered most of southeast Alaska and extended past the present coastline onto the Continental Shelf. During waxing, and to some extent waning, stages, the dominant trend of glacial erosion was controlled by northwest–southeast trending fault structures. Such structural control made possible the formation of exceptionally long fiords. However, during the time when glacial ice dominated the landscape, the flow and therefore erosional trend was seaward across the structural grain. Although today’s glaciers are only minor remnants of their Pleistocene predecessors, they remain important agents in coastal modification. Southeast Alaska, tectonically active since the early Palaeozoic, is presently being uplifted at one of the fastest rates in the world (Skelkregg 1974–1976). In places, uplift is more than 3.5 cm/year, a rate that helps account for the pres­ence of glaciomarine deposits at heights of 230 m above

sea level near Juneau. Some of this uplift is thought to be the result of rebound that occurs with deglaciation. Although general rebound is still occurring, localised rebound is partially responsible for a rapid uplift that is being recorded near retreating glaciers (Hicks and Shofnos 1965).

References Black RF (1974) Geology and ancient Aleuts, Amchitka and Umnak Islands, Aleutians. Arctic Anthropol 11:126–140 Dupre WR (1977) Yukon delta coastal processes study. University of Houston, Houston, TX Hamilton S, Shennan I (2005) Late Holocene relative sea-level changes and the earthquake deformation cycle around upper Cook Inlet, Alaska. Quat Sci Rev 24:1479–1498 Harper JR (1978) Coastal erosion rates along the Chukchi Sea coast near Barrow, Alaska. Arctic 31:428–433 Hayes NO, Ruby CH, Stephen MF, Wilson SJ (1976) Geomorphology of the southern coast of Alaska. In: Proceedings of the 15th Coastal Engineering Conference, Honolulu, HI, pp 1992–2008 Hicks SD, Shofnos W (1965) The determination of land emergence from sea level observations in southeast Alaska. J Geophys Res 70:3315–3320 Jorgenson MT, Brown J (2005) Classification of the Alaskan Beaufort sea coast and estimation of sediment and carbon inputs from coastal erosion. Geo Marine 492:32–45 Molnia BF (1986) Glacial history of the northeastern Gulf of Alaska-a synthesis. In: Hamilton TD et al (eds) Glaciation in Alaska - The geologic record. Alaska Geological Society, pp 219–236 Selkregg LL (ed) (1974–1976) Alaska regional profiles, 6 vols. University of Alaska, Anchorage, AL Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw-Hill Book Company, New York

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Trabant DC, Krimmel RM, Echelmeyer KA, Zirnheld SL, Elsberg DH (2003) The slow advance of a calving glacier; Hubbard Glacier, Alaska, U.S.A. Ann Glaciol 36:45–50 Wahrhaftig C (1965) Physiographic divisions of Alaska. United States Geological Survey Professional Paper 482, Washington, DC Waller RM (1966) Effects of the earthquake of March 27, 1964, in the Homer Area, Alaska. United States Geological Survey Professional Paper 542-D, Washington, DC

Walker HJ (1974) The Colville River and the Beaufort Sea: some inter­ actions. In: Reed JC, Sater JE (eds) The coast and shelf of the Beaufort Sea. The Arctic Institute of North America, Washington, DC, pp 513–540 Walker HJ (1991) Bluff erosion at Barrow and Wainwright, Arctic Alaska. Z Geomorphol 81:53–61

1.2  Washington Maurice Schwartz · Thomas Terich

1.  Introduction The coastline of the state of Washington is 4,296 km long and consists of three segments: the eastern coast and islands from the Canadian border in the north down to the straits and islands of Puget Sound in the south, the northern coast of the Olympic Peninsula along the Strait of Juan de Fuca, and the outer, western, coast facing the Pacific Ocean. Along the eastern side of the Olympic Peninsula is the Hood Canal, a 104 km long inundated glacial trough. An account of the coastal features of Washington State was provided by Shepard and Wanless (1971) and a geomorphic classification of the ocean coast by Terich and Schwartz (1981). Locations of places mentioned can be found in the Washington Atlas and Gazetteer published by the DeLorme Mapping Company, P.O. Box 298, Freeport, Maine 04032, USA http://www.delorme.com. Further topographic de­­ tails are shown on the relevant 1:24,000 Washington 7.5 min maps produced by the United States Geological Survey, Denver, Colorado 80225, USA Reference can be made to the Washington State Department of Ecology collection of air oblique photographs of the coast at http://apps.ecy. wa.gov/shorephotos/. The geology of the Washington State coast is dominated by the folded and faulted Tertiary formations of the Olympic Peninsula and the extensive Pleistocene glacial drift and associated deposits of the Puget Sound region, which extend northward across the Canadian ­border and also occupy valleys and mantle coastal slopes along the Pacific seaboard. There are Mesozoic and Palae­ ozoic formations south of Bellingham and in the San Juan Islands. There was intensive Pleistocene glaciation, producing rugged mountains and wide mantles of glacial drift deposits in the lowlands. Shepard and Wanless (1971) commented that although Ozette Lake (12.8 km long and over 90 m deep) lies in a basin scoured by an ice sheet, the coast between Cape Flattery and Quillayute River shows little evidence of glaciation: there are no fiords. They ­suggested that the ice was relatively stagnant in the coastal fringe, where it melted to deposit a glacial drift mantle.

Cliffs and bluffs along much of the coast and around islands are cut into this glacial drift, and sand and gravel derived from these has produced beaches and spits, especially on the islands, along the eastern coast, and around Puget Sound. Bedrock outcrops, mainly of Mesozoic and Tertiary formations, are found on headlands and along shores where the glacial drift mantle has been removed by marine erosion, and are more prominent on the Pacific Ocean coast. Stacks and natural arches along the Washington coast between the Quillayute River and Point Grenville are the outcome of differential erosion of Eocene (Quinault Formation) volcanic rocks, siltstones and sandstones and the Miocene Hoh Rock Assemblage, which consists of a basic mélange of chaotically mixed blocks of hard sandstone and basalt in a matrix of softer mudrocks and an overlying steeply tilted and overturned formation of sandstone, siltstone and conglomerate. Evidence that the Hoh mélange rose discordantly as diapirs or piercement structures through the overlying sedimentary beds can be seen in sea cliff exposures where the mélange is bounded on both sides by strata steeply dipping outward; and in offshore sub-bottom profiles indicating a doming of otherwise horizontal beds. Due to the easily eroded soft claystone and broken siltstone within it, the Hoh mélange forms embayments along the coast between headlands of the more resistant sedimentary beds of the upper Hoh Rock Assemblage and the Quinalt Formation. Some large blocks of mélange hard sandstone and basalt still remain as erosional remnants in stacks offshore. The coastal climate is cool temperate and humid, with extensive winter snow in the mountains. The Pacific Coast region has an average annual rainfall of about 2,000 mm increasing to 2,875 mm in the rain forest area inland of La Push. Along the forested coastal strip the principal trees are Western hemlock (Tsuga heterophylla), western arborvitae (Thuja plicata), Douglas fir (Pseudotsuga menziesii), Sitka spruce (Picea sitchensis), and Lodgepole pine (Pinus contorta). Westerly winds prevail, but easterlies are frequent, particularly in winter. On the Pacific coast onshore winds are highly variable, between SSW and NNW. East of  the Olympic Mountains rainfall is reduced in a rain shadow

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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region, the Sequim district being well known for its dry climate (annual rainfall down to about 400 mm), but the eastern coast is also relatively wet (1,000–1,200 mm). Tides are typically mesotidal, from about 2 m at neaps to 4 m at springs. Mean spring tide ranges on the eastern coast at Bellingham and Anacortes are 2.2 m, at Seattle 3.1 m and rising to 4.0 m at Olympic in Puget Sound, but on the Strait of Juan de Fuca they diminish to 1.8 m at Port Angeles, and on the Pacific Ocean coast are 2.2 m at Cape Flattery and 2.1 m at Willapa Bay. Never­ theless, tidal currents can be strong through narrow straits, as in Deception Pass, between Fidalgo Island and Whidbey Island, where they attain 18.5 km/h (10 knots), while currents of about 3–7 km/h (1.6–3.8 knots) are measured off Cape Flattery. Wave action along the eastern coast and in Puget Sound is determined by local winds across often short fetches in straits and between islands, but on the Pacific coast westerly ocean swell accompanies locally generated and occasional storm waves. Net longshore drifting varies with configuration (Schwartz et al. 1985): Sandy Point, north of Bellingham, has grown southward but Ediz Hook and Dungeness Spit on the Strait of Juan de Fuca have grown eastward as the result of north-westerly wave action, and the latter has also been shaped by waves from the east and north east. On the Pacific coast the prevailing south-westerly waves generate net longshore drifting northward, as can be seen in the development in that direction of the Long Beach Peninsula and the northerly deflection of the mouths of such rivers as the Copalis, Moclips, Quinault, Queets, and Hoh. There are occasional local reversals, as at La Push, where wave refraction around two large stacks has caused a spit at the mouth of the Quillayute River to grow towards the south. Occasionally tsunamis generated by tectonic disturbances along the nearby plate boundary have produced giant waves on the Pacific coast. The Cascade Subduction Zone is located 85 km offshore, where the Juan de Fuca Plate has been subducting under the North American plate during the last 7,500 years. There is stratigraphic evidence of rhythmic estuarine deposits along the shores of Willapa Bay and Grays Harbour, interpreted as indicating up to six major earthquakes (magnitude 8 or 9) along this zone in the last 3,500 years, with an average return period of 500–540 years (Atwater and Hemphill 1997). At Willapa Bay serial exposures of buried forest and marsh soils, with the uppermost soil mantled by ­tsunami-deposited sand, date the most recent tsunami event as occurring in 1700 (Atwater 1996). Cor­roborating this event are historical records in Japan of a devastating tsunami reaching the coast there on January 26, 1700

(Atwater et al. 2005). This predates European settlement of the Washington coastal area, but in the 1800s coastal native people described nearby Neah Bay as being flooded in the “not very remote past without any swell”, while canoes were lifted into trees and many people were drowned. The Washington coast has been subject to changing land and sea levels, with tectonic uplift in the Olympic peninsula and along the Pacific coast associated with movements on the nearby subducting Juan de Fuca Plate (Atwater 1996) and Quaternary isostatic movements related to land depression by glacial ice loads and recovery and after deglaciations as well as oscillations of sea level. In consequence there are terraces at various levels in the coastal region, active emergence along the shores of Juan de Fuca Strait and the eastern coast and islands, and active submergence south of Cape Flattery and in the southern part of Puget Sound. Cliffs have been cut into glacial drift and bedrock formations on coastal sectors exposed to strong wave action, especially on the Pacific coast, where there are also forested bluffs, usually with some basal cliffing, and these show occasional recession by slumping and landslides, especially during storms and periods of heavy rainfall. The Pacific coast of the Olympic Peninsula is one of the great scenic coasts of the world. Differential erosion of Tertiary continental and marine sediment and volcanics has produced sea caves, stacks, arches and high cliffs. Long term coastline recession is indicated by numerous stacks and islands in coastal waters (mostly within 2 km of the shore), and it is possible that episodes of substantial cliff recession have taken place during occasional tsunamis. Many of the stacks and small islands off the coast south from Cape Flattery are residuals of harder rock, including Eocene volcanics, sandstone and siltstone, the Miocene Hoh Rock Assemblage and the Pliocene Quinault Formation. These are found sporadically within a softer sandstone and mudstone matrix in mélange formations outcropping along the coast, and are exposed and left standing offshore as the softer rocks are removed by marine erosion. Bluffs cut in glacial drift on the islands and along the eastern coast also recede as the result of recurrent local landslides. Shore platforms are not as well developed on the coast of Washington State as on other Pacific coasts (see Australia, New Zealand, Japan), but they are found on the Cape Flattery peninsula (Bird and Schwartz 2000). They include subhorizontal high tide benches produced largely by weathering processes and seaward sloping intertidalsubtidal shore platforms cut by abrasion where waves

Washington

move sand and rock fragments to and fro. Where sediment accumulations exist along the inner edges of shore platforms, they consist mostly of coarse gravel and cobbles; sandy beaches are to be found only in the occasional embayments along this stretch. Similar features are seen on some headlands along the Pacific coast and the Strait of Juan de Fuca, but shore platforms are poorly developed on the eastern coast and around the islands. On the eastern coast and island shores there are sand and gravel beaches, often with cobbles and boulders, derived from glacial drift cliffs and the sea floor. On the shores of the Strait of Juan de Fuca and along the Pacific coast sand and gravel is also supplied by rivers, and heavy driftwood concentrations are found along the shore. As on other coasts dominated by glacial drift deposits (Denmark, for example) there are numerous spits, some straight, some recurved and some cuspate. Dungeness Spit and Ediz Hook, respectively 8 and 5.6 km long, are two large spits on the southern side of the Strait of Juan de Fuca. Damming of a river and barricading the foot of a cliff, both west and updrift of Ediz Hook has caused considerable erosion at the base of this highly industrialised spit. The U.S. Army Corps of Engineers has recently initiated a maintenance program consisting of foreshore revetment and periodic beach nourishment. Smaller spits have formed on island and peninsula shores, particularly at points where the coastline changes direction (Finlayson 2006). On Cape Flattery beaches are confined to coves, but southward along the Pacific coast there are long sand and gravel beaches between rocky headlands. South of Moclips the coast consists of broad, prograded sandy beaches fronting dunes. Provenance of the accreted sand along this sector has been attributed to the Columbia River to the south, whence sand has been carried northward by longshore drifting. With the construction of large dams on the Columbia during the twentieth century, it is speculated that the progradation regime may have ceased and possibly will turn to one of erosion. The beach that prograded alongside the breakwater built on the northern side of the Columbia River mouth at Cape Disappointment is now eroding, and erosion has followed earlier progradation alongside breakwaters at Point Brown and Westport, north and south of the entrance to Grays Harbour. Coastal dunes are best developed on the Long Beach Peninsula, where there are multiple rows of beach and dune ridges, each of which was built parallel to the prograding sandy coastline. Locally there are blowouts and small transgressive dunes, as on Leadbetter Point at the northern end of North Beach Peninsula, but the extensive dune fields of the Oregon coast have not developed here.

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Similar dune morphology has developed along the coast south of Copalis Beach to Ocean Shores and Point Brown, and between Westport and Cape Shoalwater. Otherwise there are only minor dunes, chiefly behind beaches near the mouths of rivers and on some of the spits in northern Washington. One small dune at Fort Worden State Park, near Port Townsend, has been planted with dune grass to maintain it in an intensively used area. The southern part of the Pacific coast is backed by a gently rolling plain, with two large estuarine lagoons, Willapa Bay and Grays Harbour, developed by the submergence of the mouths of valleys cut by rivers during phases of lower sea level, followed by the growth of barrier spits on the ocean shore, constricting their entrances. Northward migration of the main Willapa Bay inlet channel is believed to be the cause of severe beach erosion on the northern shore, at Cape Shoalwater, where the coastline has receded up to 3.7 km over a 76 year period. Reference has been made to the breakwaters built to stabilise the entrance to Grays Harbour. There are small estuaries at the mouths of rivers on the Pacific coast, notably the Quillayute and Queets, and the Pysht River on the Strait of Juan de Fuca, and the growth of the compound Dungeness Spit has almost enclosed the waters of Dungeness Bay as a coastal lagoon. Crockett Lake on Whidbey Island is an example of a coastal lagoon enclosed by a wave-built barrier of sand and gravel. Several deltas have been built on the eastern coast where rivers have deposited large loads of sediment brought down from mountainous catchments. They include the deltas of the Lummi and Nooksack, the Samish, Skagit and Stillaguamish deltas, the Snohomish delta near Everett, the much modified Puyallup delta (largely converted to Tacoma docks) and the Nisqually delta at the southern end of Puget Sound. The Skohomish River has built a marshy delta at the Great Bend on Hood Canal. On the Strait of Juan de Fuca coast the Elwha delta is prominent. Marshes and swamps (sloughs) have formed on the more sheltered parts of estuary and lagoon shores, and in straits and bays protected from strong wave action by ­peninsulas, islands or spits. There are corridors of fresh water swamp between the north-south dune ridges on Long Beach Peninsula, some of which are used to cultivate cranberries.

2. The Coast of Washington from North to South Point Roberts is a rectangular peninsula, an isolated segment of the United States, cut off because the Canadian

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boundary, the 49th parallel of latitude, runs westward across Boundary Bay to intersect it between Maple Beach and Boundary Bluff. It has beaches of sand and gravel bordered by multiple sand bars. Drayton Harbour is a NW-facing horseshoe-shaped bay bordered on the western side by steep wooded bluffs, and partly enclosed as a lagoon by the long Semiahmoo spit that has grown north-east toward Tongue Point opposite the harbour structure at the town of Blaine. This spit shelters salt marshes and mudflats exposed at low tide in Drayton Harbour on its landward (eastern) side, while the more exposed western shore consists of a beach of well-rounded pebbles and sand that continues southward below wooded bluffs and cliffs 20–25 m high, with occasional landslides in grey clayey sand (glacial drift). At the southern end of Semiahmoo spit, adjacent to the bluff and beside a sewage treatment site, is a large Indian midden over 6 m thick, dating from 4,000 bp onwards and rather overgrown. The beach passes in front of slumping bluffs south towards Birch Point (>Fig. 1.2.1). Birch Bay is similar in outline to Drayton Harbour, but faces SW and has no enclosing spits. Mean spring tide range is about 3.6 m, and at low tide extensive sand bars are exposed in front of a gravel beach segmented by groynes. The beach outline is an asymmetrical (log-spiral) curve, produced by the refraction of waves from the NW, and although there are episodes of northward and southward drifting net longshore drift is zero. A succession of emerged coastlines in the hinterland of Birch Bay, marked

by a stairway of glacial rebound terraces, follows the same asymmetrical curve. South of Neptune Beach is Sandy Point, a large spit that has grown southward as the result of wave action sweeping sand and gravel along the coast (>Fig. 1.2.2). In the 1960s canals were dredged within the spit to form a marina, and there are now many houses behind the beach and around the boat channels. Unfortunately, beach erosion has become a serious problem, and residents have built sea walls and boulder ramparts on the shore to protect their properties. The spit is low-lying, and occasionally overwashed and flooded during storms. In the 1920s it was temporarily breached, and the storm of Easter 1977 swept sand and gravel in from the beach and damaged several houses. Within Lummi Bay, to the south, are the eroded (and now embanked) remains of the Lummi delta, which decayed after the diversion of the Nooksack River eastward into Bellingham Bay in 1877 (see below). Portage Point is at the south-eastern end of the Lummi Nation Indian Reservation peninsula. The Portage (>Fig. 1.2.3) is a tidal divide (intertidal tombolo) where a road is exposed at low tide, giving access to Portage Island. Brant Point is a sandy spit with shoals curving out in the lee of Portage Island into Bellingham Bay. There is a cliff cut in glacial drift on the southern end of Portage Island at Point Frances (>Fig. 1.2.4) which is typical of the intermittently active cliffs in this region. On the northern side of Bellingham Bay the Nooksack River has a growing delta, and the bay water is discoloured

⊡⊡ Fig. 1.2.1 Slumping cliff in glacial drift north of Birch Point. (Courtesy Geostudies.)

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⊡⊡ Fig. 1.2.2 Air view of Neptune Beach and the spit that has grown southward at Sandy Point. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.3 The intertidal tombolo linking Portage Island to the Lummi Peninsula. (Courtesy Washington Department of Ecology.)

by brown silt and clay when the river floods. The Nooksack River has a northern branch, Lummi River, which flows to the sea in Lummi Bay on the northern side of the peninsula. In the mid-nineteenth century the Nooksack had become choked with a massive log jam just below the divergence of the two rivers, and some of the water and sediment discharge was carried down Lummi River into Lummi Bay. In 1876–1877 the log jam was removed, and the Lummi River diked off. After that the Nooksack delta began to grow out into Bellingham Bay (Shepard and Wanless 1971). The San Juan Islands are scattered, mainly forested islands of various sizes and shapes in the southern part of the Strait of Georgia. They represent the partial marine submergence of a mountain range, dissected by glaciated

valleys that are now fiords. San Juan Island is the largest of the group, about 29 km long with an average width of 10 km, a hilly island rising to the grassy granite slopes of Mount Young, 198 m high. It has steep coasts and rocky shores, declining to a narrow peninsula in the south-east, extending to Cattle Point. On the southern coast False Bay is a shallow rounded embayment with a wide intertidal zone. To the north of San Juan Island is Waldron Island, consisting of a south-eastern upland of Upper Cretaceous shaly sandstones and conglomerate culminating southward in the steep-sided narrow Point Disney, which has vertical cliffs up to 150 m high on its western flank. On the eastern coast is rocky Mail Bay, and to the north a lowland of glacial drift rising to 30 m above sea level. Stratified glacial

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⊡⊡ Fig. 1.2.4 The cliff at Point Frances on the southern coast of Portage Island. (Courtesy Washington Department of Ecology.)

deposits outcrop in 30 m cliffs near Point Hammond at the north-eastern end of the island, and there are several small promontories of exposed bedrock, the most prominent at Fishery Point in the north-west. From here North Bay curves out to a sharp cuspate point of glacial drift projecting westward to Sandy Point. To the south Cowlitz Bay is another curved bay that extends south to Point Disney, and near the south-eastern end is intertidal Mouatt Reef, rising a metre above low tide. To the north of San Juan Island is Spieden Island, a narrow ridge rising to 116 m. Nearby is Flattop Island, which has a steep rocky north-west coast and a surface that is not flat, but slopes south-east with the dip of the rock formations. Gull Rock, off the north-west point of the island, is split by a chasm cut out of shale between hard conglomerates. The hard rocks form reefs, notably White Rocks, one of which rises 10 m above high tide level. To the west is Stuart Island, almost split by elongated Reid Harbour, with Turn Point at its western end looking across Haro Strait to the similar Canadian islands that border Vancouver Island. Also to the north of Waldron Island is Bare Island, a small island and seabird colony rising 12 m above sea level and almost unvegetated. Narrow inlets have been cut in steeply dipping shales on the east and west coasts, and there is a cap of glacial drift as well as much guano. Orcas Island is M-shaped, with a smooth ­north-eastern coast following the geological strike and to the south

s­ teep-sided ridges separated by deep inlets, East Sound, West Sound and Deer Harbour. Mount Constitution, the highest peak in the San Juan Islands, rises to 759 m. Point Doughty is low and narrow, projecting into President Channel, and Freeman Island is a small island to the south that has a broad flat rock platform exposed at low tide. There are broad flat rocky reefs off the north coast, including Parker Reef, exposed at low tide. Lopez Island is a lowlying island to the south of Orcas Island. On the mainland coast Bellingham Bay is a broad bay opening southward and bordered by steep bluffs and occasional cliffs. Relics of the wooden supports for the Bellingham and British Columbia Railroad stand offshore along the SE coast, and at Bellingham the Georgia Pacific water treatment lagoon is prominent on the shore. Aerating sprinklers and introduced bacteria and fungi were used to break down chemicals generated by the paper works and discharged into the sea, but this reduction of pollution in Belling­­­ ham  Bay resulted in the return of the shipworm Teredo, which bored into the wooden pilings that support piers and attacked the rafts of logs that used to be floated in by sea. Fairhaven is a historic neighborhood, dating from the mid-nineteenth century. The coast has been modified by reclamation. Towards the end of the nineteenth century large quantities of sand and gravel were sluiced from Poe’s (Post) Hill Point, which was an 18 m high promontory, and deposited in the bay to form new land that was then called Commercial Point for the development of

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shipbuilding yards, boat piers and factories. Depletion of the beach at Marine Park, Fairhaven, was countered in 2004 by emplacing an artificial beach of shingle. South of Fairhaven Chuckanut Mountain (>Fig. 1.2.5) descends to a steep forested coast on Cretaceous sandstones, passing southward across the Oyster Creek Fault to Palaeozoic metasediments. Below the steep slopes mudflats are exposed at low tide on the shores of Samish Bay. The view south from Chuckanut Mountain is of a wide coastal plain which includes the Samish River delta (>Fig. 1.2.6). ⊡⊡ Fig. 1.2.5 Steep coast on Cretaceous sandstones in Larrabee State Park, below Chuckanut Mountain south of Fairhaven. Honeycomb weathering is well developed on these feldspathic sandstones (Mustoe 1982). (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.6 Coastal plain formed by reclamation of swamps and intertidal areas on the Samish delta. (Courtesy Geostudies.)

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Samish Island, a wooded flat-topped ridge (two former islands linked by a sandy tombolo), was attached to the mainland by reclamation of intertidal land between earthen dikes by Dutchmen more than a century ago. Offshore is a group of islands including high Cypress Island and low-lying Guemes Island. Sand and gravel eroded from cliffs has supplied beaches and spits shaped by wave action on the shore of Lopez Island (>Fig. 1.2.7). South of the Guemes Channel is Fidalgo Island, the inner part of which is separated from the mainland only

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by a marshy isthmus followed by the Swinomish Channel. Shannon Point is a cliffed promontory at the north-­western end with glacial drift (Vashon drift (c. 13,000–15,000 years bp.) of the Fraser glaciation) over strongly folded and metamorphosed Mesozoic sediment and volcanics and a crystalline basement of Devonian rocks. Fidalgo Head on the west coast is a grassy bluff with outcrops of serpentine rock, above the swirling waters of Burrows Channel, looking across to the San Juan Islands. To the south is Deception Pass (>Fig. 1.2.8).

Deception Pass is located between Fidalgo Island and Whidbey Island and is a deep and narrow strait through which the tides swirl at up to 18.5 km/h (10 knots). It is bordered by steep rocky cliffs. When it was discovered by Vancouver in 1792 he thought it would lead through to a harbour, but when he found otherwise he named it from having been deceived. West Point (>Fig. 1.2.9) has been truncated by strong tidal currents. Whidbey Island is about 145 km long, parallel to the mainland coast and extending southward to Seattle. On the ⊡⊡ Fig. 1.2.7 Cuspate spit, Flat Point, Lopez Island, San Juan Islands, with Shaw Island on the right. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.8 Deception Pass, showing strong tidal current. (Courtesy Geostudies.)

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west coast high cliffs of sandy glacial drift face the Strait of Juan de Fuca, with beaches and a looped barrier (>Fig.  1.2.10). Point Partridge and Admiralty Head are prominent headlands, and on the latter is Fort Casey, where an old searchlight tramway protrudes 4  m from the cliff below an artillery fort, indicating recent recession of the cliff. This part of Whidbey Island has the Weird Pits, which are kettle holes up to 30 m deep which formed as the result of collapse following the melting of chunks of ice left in the glacial drift after the glaciers receded. Near Point Partridge ⊡⊡ Fig. 1.2.9 West Point, Deception Pass, with a strong ebb current flowing past the rocky outcrop and driftwood on the beach. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.10 Ebey’s Landing on the west coast of Whidbey Island has a looped barrier of sand and gravel enclosing Parego’s Lagoon, backed by bluffs 80 m high. (Courtesy Geostudies.)

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one of these has been excavated by the sea to form a cove with a little lagoon. On the east coast of Whidbey Island a succession of promontories separates coves and harbour bays. Bluffs with a scrub or forest cover alternate with cliffs exposing glacial drift deposits, and there is occasional slumping in the bluffs. At low tide a gravelly foreshore is exposed at the village of Coupeville, founded in 1852 on slopes descending into Penn Cove. Skagit River has built a delta on which the river splits into North Fork and South Fork, on either

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side of Fir Island, and there are relics of many distributary channels. Camano Island, parallel to Whidbey Island east of Saratoga Passage, has wooded bluffs and occasional cliffs. Camano Head (100 m) is one of the tallest, steeply sloping bluffs in the archipelago, with tanglewood scrub sliding down the slope. South of Stanwood the Stillaguamish River has built a delta into Port Susan, a bay sheltered from westerly winds and waves by bluffs in the lee of Camano Island. The coast southward along the Tulalip Indian Reservation has wooded bluffs and sand and gravel beaches, with some minor spits. Mission Point is a hooked sand spit sheltering Tulalip Bay. A delta has been built where the Snohomish River and Steamboat Slough curve out on either side of Smith Island, north of the town of Everett. This town is built on low hills beside the valley of the Snohomish River. The waterfront harbour on the shores of Port Gardner Bay is sheltered by a chain of low islands extending north to Jetty Island, partly artificially formed by the dumping of spoil from dredging. Reclamation of the Everett waterfront has formed a series of docks and marinas and a large U.S. naval station at the southern end. Puget Sound was cut by glaciers and meltwaters which formed deep, steep-sided channels, into which the sea flooded to form marine inlets that wind and branch deep into the state of Washington; inlets bordered by forested bluffs. The islands are generally steep-sided and flat or gently undulating, possibly as the result of marine

­ lanation of the glacial drift deposits at higher sea levels. p The rise and fall of the tides in Puget Sound is about 2.1 m in the north, increasing to nearly 3 m in the south, and at low tide sandy and gravelly beaches and muddy fringes are exposed. Spits are found on many of the islands (>Fig. 1.2.11), and there are intertidal and subtidal shoals. South from Everett bluffs and cliffs form the eastern coast of Possession Sound, and extend past Mukilteo to Norma Beach, where the coastal suburbs of Seattle begin. Bluffs and small headlands with some coves and beaches line the Seattle coast. Seattle has spread across the isthmus of ridges and valleys that trend south to north on the eastern side of Puget Sound, has a coastline of wooded bluffs with headlands and bays, and some of which have been reclaimed for port development. West Point is a cuspate spit that carries a City of Seattle sewage treatment plant, a small Metro lighthouse marking the northern point of Elliot Bay, and a Coast Guard building. Elliot Bay provided a good harbour in deep water, below a steep wooded hillside with ravines and springs. This part of the coast has been much modified by land reclamation; in 1902–1930 Denny Hill was removed and the earth and rock sluiced from it was dumped to reclaim former tideflats in Elliot Bay. Tacoma is a large city on a wide plateau near the southern end of Puget Sound. Commencement Bay is broad, with the delta of Puyallop River at its head, modified to form Tacoma docks. The coast southward is dominated by ⊡⊡ Fig. 1.2.11 Trailing spit extending south east from Smith Island in the Strait of Juan de Fuca. (Courtesy Geostudies.)

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steep bluffs cut in glacial drift, extending past the narrow promontory at Point Defiance. Towards the southern end of Puget Sound the Nisqually River has built a shrubby deltaic plain with extensive marshes. The head of Budd Inlet, to the west, has been enclosed to form Capitol Lake, alongside Olympia. Bainbridge Island, on the western side of Puget Sound opposite Seattle, is a broad plateau of glacial drift bordered by bluffs, separated from the mainland by Agate Passage. Restoration Point has an emerged shore platform 6 m above the modern one, probably uplifted tectonically. Point Monroe has a looped sand barrier enclosing a lagoon at the northern end of Bainbridge Island. To the south is Vashon Island, similar to Bainbridge Island, attached to Maury Island by way of an isthmus at Portage. On the west coast of Puget Sound opposite the cliffy southern end of Whidbey Island is Point No Point, a rounded or lobate spit like many others on Puget Sound, shaped by waves delivering sand and gravel but blunted by strong tidal currents. Hood Canal, the long channel that diverges from Admiralty Inlet near Port Ludlow is generally 3–5 km wide as it runs southward past Hoodsport to the Great Bend, and 1.5–3 km wide as it swings north-east to its marshy head at Lynch Cove, near Belfair; it is bordered by steep wooded slopes that are sometimes disrupted by landslides, the scars of which persist for some years. Inflowing streams have built small deltas, such as that of the Dosewallips River, and a large delta has been built by the Skokomish River south of the Great Bend near Union. Longbranch Peninsula, between Case Inlet and Carr Inlet, has springs flowing from cliffs of gravel and clay on its west coast. The Hood Canal opens northward into Admiralty Inlet near Port Ludlow, sited on a slight hill overlooking a harbour. To the north, Marrowstone Island and Indian Island are islands of glacial drift that run north-south. They are fringed by bluffs and minor cliffs, and have beaches of sand and gravel and spits, such as the cuspate spit at Marrowstone Point. Port Townsend on the east coast of the Quimper Peninsula, looks across Admiralty Inlet to Whidbey Island, and has bluffs fronted by a reclaimed shore area (Zelo et al. 2000). Beach sand drifts northward past Fort Worden. Point Wilson at the northern end of Quimper Pen­ insula, on the western side of the entrance to Admiralty Inlet and Puget Sound, is a sand and gravel spit with low dunes. It has been shaped by waves arriving from the south-east, across Admiralty Inlet, and from the west, along the Strait of Juan de Fuca. The Quimper Peninsula, bordered to the west by Discovery Bay is broad and flat-topped with steep forested

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bluffs. At Fairmont, a village at the head of Discovery Bay, salt marshes and mudflats are exposed as the tide falls. Offshore is Protection Island, which became a national wildlife refuge in 1982. West of Discovery Bay is the broad Miller Peninsula, and then Sequim Bay, which is essentially a lagoon almost enclosed by Gibson Spit (>Fig.  1.2.12) which has grown southward on the western side and Travis spit, which has grown eastward. Strong tidal currents flow through the gap. Dungeness Spit (>Fig. 1.2.13) is a sand and gravel spit nearly 10 km long, a long narrow ridge of sand and gravel, piled with driftwood, with lobes of washed-over sand, runs out eastward to the distant lighthouse (Schwartz et al. 1987). Sediment from eroding cliffs in glacial drift to the west has drifted alongshore to build the spit, which is still growing eastward. It is backed by a lagoon, Dunge­ness Bay, containing truncated Cline spit and broader Graveyard Spit, running sharply back from the middle of Dungeness Spit, and intervening marshes and lagoons (>Fig. 1.2.14). Dungeness Spit has been supplied with sand and gravel from eroding cliffs in glacial drift to the west. The steep cliffs cut in sand and gravel between Dungeness Spit and Port Angeles are receding, and pose a problem for houses built on cliff-top land. Cliff recession occurs as the result of basal wave attack, but there is also slumping, particularly after wet weather when the glacial drift deposits become overloaded with groundwater and are unstable. Port Angeles stands in the lee of the Ediz Hook spit. Ediz Hook is a spit 5.6 km long that has grown out from the coast east of the Elwha delta, supplied with sand and gravel from eroding cliffs and from the Elwha River. In recent years erosion has become severe along the shores of this spit, partly because of the reduction of sand and gravel supply from cliffs that have been stabilised and partly because of the damming of the Elwha River, so that the shore is now heavily armored with large rock boulders, which preserve the spit artificially as a breakwater for Port Angeles. The Elwha River is incised in a deep valley. It has built a delta, but this is eroding because dams have intercepted most of the sediment previously brought down the river to the coast. West from the mouth of Elwha River the coast consists of bluffs and cliffs, interrupted at the mouths of many small river valleys. Pysht is at the mouth of two valleys, and to the west there is a rocky headland and a prominent sea stack to the west on Pillar Point. Clallam Bay is a village with a small fishing harbour, where a walkway leads out from the coastguard station to a steep headland and Sekiu is a village on the western shore of Clallam Bay, with a sheltered harbour. To the west are cliffs fronted by shore platforms exposed at low tide.

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⊡⊡ Fig. 1.2.12 Gibson Spit (left) and Lagoon at Washington Harbour and Travis Spit (right) on Miller Peninsula constrict the entrance to Sequim Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.13 Dungeness Spit. (Courtesy Journal of Coastal Research.)

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⊡⊡ Fig. 1.2.14 The eastern end of Dungeness Spit, showing the broad Graveyard Spit and the narrower Cline Spit built by easterly wave action. Directions of longshore drifting are indicated. Strong tidal currents flow through the constricted passage in and out of Dungeness Bay, to the left: they attain 56 cm/s as the tide rises and 60 cm/s (about 1 knot) as it ebbs, and the channel floor has a gravel lag deposit, in contrast with mud under the lagoon and sand near the shores. Dungeness River, on the right, reaches the sea by way of an outlet deflected westward by longshore spit growth. (Courtesy Geostudies.)

⊡⊡ Fig. 1.2.15 Old Hat island at Seal Rocks, east of Neah Bay, Strait of Juan de Fuca. (Courtesy Geostudies.)

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Seal Rocks, east of Neah Bay, include high Old Hat islands (>Fig.  1.2.15), surrounded by subhorizontal high tide shore platforms produced largely by weathering processes (Bird and Schwartz 2000). Neah Bay, to the east of Cape Flattery, has a harbour protected by a breakwater linked to Waadah Island, and has a crescent-shaped beach of pure white sand connecting Bahada Point. Tatoosh Island (>Fig.  1.2.16) is a flat-topped cliffedged island offshore, the northwesternmost point of the 48 coterminous United States at the western end of the Strait of Juan de Fuca. Cape Flattery has massive sandstone cliffs with basal rock ledges, looking out to Tatoosh Island. The peninsula consists of Lower Tertiary (mainly Eocene) stratified sandstones and siltstones and, locally massive or conglomeratic. The cliffs show weathering features, including tafoni and caves, structural ledges, and basal subhorizontal high tide shore platforms (Bird and Schwartz 2000). High cliffs run south from Cape Flattery towards Waatch Point, where there are wide rocky shore platforms cut by abrasion as waves move sand and gravel to and fro. Hobuck Beach at the mouth of Waatch River curves round Mukkaw Bay towards Anderson Point. Sooes River flows northward behind a barrier spit to a deflected outlet. Anderson Point has subhorizontal high tide shore platforms around islands and rock stacks. Shishi Beach, 4.8 km of sandy shore, runs south from Portage Head in the Makah Indian Reservation; there are numerous grey

jagged reefs and stacks offshore. Point of the Arches is a cliffy headland with caves and tunnels and many stacks offshore. Between Point of the Arches and Cape Alava is wilderness coast in the Olympic National Park, accessible only to walkers. Cape Alava is the most westerly point in the coterminous United States. South from Cape Alava, near Ozette, the coastline is a series of beaches a few ­hundred metres long with a headland at each end. The headlands are 30–50 m high, steep and thick with salal bush, and impassable at high tide. Alava Island stands as a mesa in the ocean, first and biggest of a group of sea stacks, reefs, isles and boulders known as The Flattery Rocks. One of the first headlands south of Cape Alava is Wedding Rock, a dark-basalt cliff with broken rocks at its base. These rocks have served as a canvas for ancient artists, with petroglyphs carved on the rocks. Sand Point, down a 15 km walking trail from Ozette, is notable for its shore pools exposed when the tide is out. Ozette Lake is elongated parallel to the coastline, and one report suggests that it was impounded by a bouldery glacial moraine on its western shore (Alt and Hyndman 1984). Its great depth (90 m) indicates that it occupies an ice-scoured depression. The coast continues southward past Point Johnson and several other rocky promontories. Hole in the Wall is on a small headland penetrated by a wave-cut cave. There are many outlying sea stacks of dark basalt. Rialto Beach, a curving bank of pebbles and cobbles backed by heaps of driftwood worn smooth and clean by the abrasive waves, runs out from the mouth of Quillayute River at La Push. The harbour is ­protected by a large ⊡⊡ Fig. 1.2.16 An air view of Tatoosh Island, looking towards Cape Flattery and showing high-tide shore platforms.

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breakwater, strewn with driftwood, linked to one of several forested flat-topped islands. To the south are First, Second and Third Beaches, extending down to rugged Teahwhit Head. 27 km of steep forested wilderness coast extends between La Push and Ruby Beach, which takes its name from the polished black stones that can be found on the shore. This part of the coast is accessible only on foot, along Olympic National Park trails that run behind prominent Hoh Head and a succession of rocky capes and curving bays, with many islands and stacks offshore (>Fig. 1.2.17). Destruction Island is a bold, steep-sided island, 1.5 km long and 5 km offshore (Shepard and Wanless 1971), one of the very many islands and stacks off the oceanic Washington coast. To the south Kalaloch Creek swirls out past a slumping meander cliff to the sea. Southward the coastline straightens, and is fringed by a beach of sand and gravel. At Queets the beaches are again numbered rather than named: First Beach, Second Beach and Third Beach. Tunnel Island is penetrated by a large cave. There have been recurrent tsunamis on this coast. Cape Elizabeth is a flat-topped promontory with cliffs cut in soft stratified siltstone and conglomerate (Quinault Formation) which run southward to the mouth of Quinault River. A 1902 photograph shows that a double arch in Eocene sandstone and conglomerate existed here, but before 1970 the inner arch collapsed. South of Taholah is Point Grenville, a rocky promontory of marine volcanic ⊡⊡ Fig. 1.2.17 Hoh Head, a cliffed promontory of hard sandstone that has persisted while adjacent sectors of weaker mélange have been cut back. (Courtesy Washington Department of Ecology.)

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rocks, rising abruptly from the ocean, into which it extends a short distance in a semicircular shape. Outlying to the southeast are two pyramidal rocks 20–25 m high. Moclips is a village on top of bluffs with pine and cedar forest at the mouth of Moclips River. The straight coast ­continues, with bluffs and cliffs rising behind a long sandy beach. Roosevelt Beach is a firm sandy surf beach in front of cliffs of soft dark sand. The barrier beaches continue southward to the mouth of the Columbia River. Roosevelt Beach narrows southwards past Copalis River to Copalis Beach, where the steep wooded bluffs, locally slumping, give place to dunes behind a narrow shingle beach with much driftwood. At Sampson the beach and dunes become a barrier spit on the northern side of Grays Harbour. On it are extensive seaside resorts, including Ocean Shores. Grays Harbour is a large estuarine lagoon which narrows eastward to the mouth of Chehalis River, where Aberdeen and Hoquiam are log-shipping and lumber-milling towns. At Point Brown, on the northern shore, North Jetty was built to stabilise the entrance. A wide sandy beach, backed by dunes, built up after the jetty was first built, but is now eroding away, and at Ocean Shores a large condominium block which was threatened by shore erosion and protected by massive boulder rampart now projects on to the beach. Westport on the broad sandy spit that borders the southern part of Grays Harbour stands behind the sand dunes, with long jetties built early in the twentieth century

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to stabilise the entrance to Grays Harbour and disperse the sand shoals that impeded navigation: at first the jetties led to widening of the sandy beach on either side of the entrance, but in the last few years there has been rapid erosion on the southern shore, and the Corps of Engineers has had to deposit about 600,000 cubic metres of sand and pebbles to make a wide isthmus that prevents the sea breaking through into Half Moon Bay and threatening Westport town. Cape Shoalwater is a low promontory on the northern side of the entrance to Willapa Bay, and has been subject to severe marine erosion. Comparison of historical maps shows that Cape Shoalwater was been cut back 3.7 km over a 76 year period, and that since 1967 it has been receding about 30 m/year. Since the early charts were made the entrance to Willapa Bay has widened, and the erosion is partly due to the deep-water channel impinging on the Cape Shoalwater shore (>Fig. 1.2.18). Willapa Bay, once known as Shoalwater Bay, is a large coastal lagoon with a wide shoaly entrance, an irregular outline, rocky headlands and marshy coves, and the high forested Long Island; the average tide rises and falls about 2 m. It has been calculated that Willapa Bay exchanges 12 million gallons of water per tide: on the west coast only the Columbia River, San Francisco Bay, and Puget Sound exchange more than this. Long Beach Peninsula is a barrier spit, consisting of sand brought down to the sea by the Columbia River, which has accumulated on the coast north of the river

mouth. A wide beach is backed by grassy and forested ­parallel dune ridges, with Willapa Bay on the landward side, the dunes running from south to north, and separated by elongated hollows containing lakes and bogs. Long Beach extends north from Cape Disappointment to Leadbetter Point at the entrance to Willapa Bay. South of Seaview the dunes come to an end, interrupted by cliffy North Head and the rocky Cape Disappointment promontory. From Cape Disappointment there is a view of the long North Jetty, with much driftwood on the sandy beach that has grown very wide on its northern side. The breakwater was built to stabilise the northern side of the mouth of the Columbia River, and comparison of early maps and charts with the modern configuration indicates that a wide beach has formed on its northern side. Nearby is a deep, steepsided rocky inlet. Large quantities of sand were carried down the Columbia River and swept into the sea in times of flood, and ocean waves have carried this onshore to build the beaches and dunes of North Beach to the north (and Clatsop Plains, in Oregon to the south). In recent decades the sand supply has diminished, possibly because of dam construction upstream, and beaches that were formerly built up with sand washed in from the sea floor now show  signs of erosion. The explosive eruption of Mount St. Helens in 1980 sent large quantities of sediment down the Toutle River into the Cowlitz River and so into the Columbia River at Longville. The sediment reaching the

⊡⊡ Fig. 1.2.18 Coastline erosion at Cape Shoalwater. (Courtesy Geostudies.)

Washington

Columbia River was largely fine-grained (silt and clay) rather than sand. At Longville the river shallowed overnight from, and dredging was necessary to 12 to 4 m restore navigability. Turbid water flowed downstream to the mouth of the Columbia River, but there was no sand accretion as a result of the Mount St. Helens eruption. Along the northern side of the Columbia River steep slopes descend to the estuary shore. There is a succession of bays and rounded promontories, and Grays River and a number of smaller streams flow into the Columbia estuary from the north. Upstream the estuary becomes shoaly, with marshy islands extending at intervals up to the Lewis and Clark bridge at Longview.

References Alt D, Hyndman D (1984) Roadside geology of Washington. Mountain Press Atwater BF (1987) Evidence for great Holocene earthquakes along the outer coast of Washington State. Science 236:942–944

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Atwater BF (1996) Coastal evidence for great earthquakes in western Washington. In Rogers, AM et al. United States Geological Survey Professional Paper 1560:77–90 Atwater BF, Hemphill Haley E (1997) Recurrence intervals for great earthquakes of the past 3500 years at Northeastern Willapa Bay. Washington. United States Geological Survey Professional Paper 1576 Atwater BF et al (2005) The orphan tsunami of 1700: Japanese clues to a parent earthquake in North America. United States Geological Survey Report 1707 Bird E, Schwartz M (2000) Shore platforms at Cape Flattery. Washington. Washington Geology, 28: 10–15 Finlayson DP (2006) The geomorphology of Puget sound beaches. Dissertation, School of Oceanography, University of Washington, Seattle, WA, 216 p Mustoe GE (1982) The origin of honeycomb weathering. Bull Geol Soc Am 93:108–115 Schwartz ML, Fabbri P, Wallace RS (1987) Geomorphology of Dungeness Spit, Washington, U.S.A., J Coastal Res 3:451–455 Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw-Hill, New York Terich TA, Schwartz ML (1981) A geomorphic classification of Washington State’s Pacific Coast. Shore Beach 49:21–27 Zelo I, Shipman H, Brennan J (2000) Alternative bank protection methods for Puget Sound shorelines. Washington State Department of Ecology publication #00-06-012

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1.3 Oregon

Paul D. Komar

1.  Introduction

2.  The Oregon Coast

The Pacific coast of Oregon extends for nearly 700 km from the Washington border on the Columbia River in the north, to the border with California in the south (>Fig. 1.3.1). The Oregon coast is characterised by considerable geomorphic variability, with long stretches of rocky shore and major headlands formed of resistant volcanic rocks, which isolate sections of sandy beaches ranging in lengths from small pockets to littoral cells that extend for 10 s to over 100 km distances along the coast; the longest is the 250 km beach from Florence to Coos Bay, backed by the Oregon Dunes Recreation Area, the most extensive coastal dune sheet in North America.

Important to the geology and geomorphology of the Oregon coast is its tectonic setting, being located within the zone of collision of three of Earth’s tectonic plates, the oceanic Juan de Fuca and Gorda plates, and the continental North American plate. The crust of the ocean plates is formed at the spreading ridges and is then carried eastward toward the continent, where being denser the ocean crust slides beneath the less-dense continental crust and is subducted. In most locations where plate subduction occurs major earthquakes are generated by the plates scraping together, but there has not been a subduction earthquake off the Oregon coast since the settlement by Euro-Americans. However, evidence has been found that such earthquakes occurred in the prehistoric past, evidence that includes estuarine marsh sediment buried by layers of sand, which had been transported far inland by huge tsunami waves that had accompanied each earthquake (Atwater 1987). Based on the numbers of such layers discovered by geologists along the coast, it  has been concluded that catastrophic earthquakes and  tsunami have occurred repeatedly in the past, with intervals ranging from 300 to 600 years. Carbon-14 dating of the buried marshes indicated that the most recent event occurred about 300 years ago, with its exact date having been established by Satake et al. (1996) to have been 26 January 1700, based on the arrival of the generated tsunami along the coast of Japan where it destroyed a number of villages. From the size of the tsunami waves that reached Japan, it was concluded that the earthquake must have been about magnitude 9, with the Oregon event having been comparable to the Sumatran subduction earthquake and tsunami in 2005. As the oceanic plates are being subducted beneath the continent, sediment that had accumulated on the seafloor as the plate moved slowly toward the coast is scraped off and added to the continental mass. This is the origin of the  Tertiary mudstones and siltstones that form the sea cliffs along the Oregon coast, while the rocky headlands are composed of volcanic rocks whose origins are also

⊡⊡ Fig. 1.3.1 The coast of Oregon and its principal communities.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Oregon

a­ ssociated with this tectonic setting. Nearly all of western Oregon has been formed by continental accretion of ocean sediment and a series of volcanic seamounts and islands, or by entire blocks of the sea floor. The oldest rocks found in western Oregon date back to the Palaeocene and Eocene geological periods, i.e. 40–60 million years ago. About 5 million years ago, during the Pliocene, the mountains of the Coast Range progressively emerged from the sea and western Oregon as we recognise it today came into existence. Upon emergence from the sea, the rain, rivers and ocean waves went to work to erode what formerly had been deep-sea crustal rocks, the basalts of former volcanoes, and ocean sediment now hardened into rock. These processes etched out the land, cutting away the weaker rocks, leaving behind the more resistant to form the peaks of the Coast Range and headlands along the coast. As on the coasts of Washington and northern California, there are numerous small islands and stacks in Oregon’s coastal waters, representing remnants of resistant rocks left behind as the coast retreated (>Fig. 1.3.2). The modern morphology of the coast is the product of the erosion that has taken place during the past 5 million years. Due to the tectonic setting of the Oregon coast, there are significant land elevation changes that affect the ­relative sea level, the change in the elevation of the land relative to that of the sea, which is globally rising at a rate between 1 and 2 mm/every year. The studies of the buried marsh deposits in estuaries that documented past occurrences of subduction earthquakes also supported the conclusion

that nearly all of the Oregon coast abruptly subsided during those tectonic events, often by 1–2 m (Atwater 1987). In contrast, much of the coast is presently rising, with that south of Florence being uplifted faster than the global rise in sea level. This has been documented through analyses of benchmarks used by surveyors, which the government re-surveys every few years to determine their elevation changes. South of Florence the tectonic rise of the land represents a net rate of land emergence above the sea of the order of 1 mm/every year, while along the north coast the relative sea level rise is on average −1.5 mm/every year (the negative value signifying that the global rise in sea level is faster than the change in land elevations). The consequence of this difference is that during historic times coastal erosion has been substantially less along the southern Oregon coast where the land is emerging from the sea, evident in the vegetated sea cliff at Bandon (>Fig. 1.3.2), in contrast with the significant erosion that is occurring along the northern Oregon coast (Komar and Shih 1993). The change from abrupt subsidence of the land during a subduction earthquake, compared with the progressive aseismic uplift now experienced, is interpreted in terms of the present accumulation of subduction strain between earthquake events, causing the slow rise of the land while the plates are locked together; the sudden release of that strain at the time of an earthquake results in the immediate subsidence of the land. The long-term net tectonic uplift of the Oregon coast, together with the cycles of sea level rise and fall during the

⊡⊡ Fig. 1.3.2 The sea cliff and offshore stacks at Bandon on the southern Oregon coast. The numerous stacks are interpreted as evidence for the rapid erosion of this shore following the 1700 earthquake when this area abruptly subsided, while the near absence of erosion today is attributed to the ongoing tectonic uplift of the land, which exceeds the rate of rise in sea level.

Oregon

Ice Ages, has given rise to marine terraces that in places form stairways up the flanks of the Coast Range. Along the south coast the highest and oldest terrace in the series reaches elevations up to nearly 500 m. The lowermost terrace, extending along much of the coast, dates back about 80,000 years. The Pleistocene terrace sands (former beaches and dunes), together with the ­underlying Tertiary mudstones and siltstones, are being eroded by the waves, forming sea cliffs that back sand beaches. Many of Oregon’s coastal communities are situated on this nearly level terrace; cities such as Cannon Beach, Lincoln City and Newport have suffered property losses as the cliffs progressively retreated (>Fig. 1.3.3). State Parks have also been impacted, with the loss of picnic grounds and camping facilities. In total, sea cliff erosion and property losses affect hundreds of km of the Oregon coast (Komar 1997). Landslides are common along the Oregon coast. The sea cliffs cut into the marine terraces are particularly susceptible to mass movement on various scales, with the occasional formation of large landslides (Komar 2004). This susceptibility is due in part to the consistency of the cliffs, landsliding being most active and destructive where the cliffs are composed of Tertiary mudstones and with their layers dipping toward the ocean. Most destructive in terms of losses of private property has been the Jump-Off Joe Landside in Newport (Komar 1997). Its initial movement began during the winter of 1942–43, affecting about 15 acres along with the loss of 15 homes. An attempt was made to re-develop the site in 1980, with plans to build

⊡⊡ Fig. 1.3.3 Sea cliff erosion at Gleneden Beach south of Lincoln City, the cliff being composed of uplifted Pleistocene terrace sands, with their erosion being the main source of sand to the beach.

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condominiums on the down-drop block of the landslide. This plan was prevented when the State rejected the developer’s request to construct a sea wall along the toe of the slide, to prevent further wave ­erosion and its continued seaward movement. The developer instead constructed the condominiums on a small remnant of the marine ­terrace to the immediate north of the Jump-Off Joe Landslide, beyond which was a second older landslide of comparable size. As the condominiums approached completion in 1981, slippage in the remnant terrace undermined the foundation, leading to their des­truction (>Fig. 1.3.4). With the developer having gone bankrupt, the  city had to cover the expenses of tearing down the condominiums. Landslides are also found on the basaltic headlands, or more precisely within the loose debris shed from the headlands that had accumulated along their flanks. Massive landslides associated with headlands have affected a few private homes, but in particular park lands such as Ecola State Park on Tillamook Head north of Cannon Beach. A number of large inactive landslides are found along the coast, believed to have formed at the time of the 1700 ­subduction earthquake; a few experienced renewed movement when their forest cover was removed by commercial logging. The differential rates of erosion and long-term retreat of the different rock types found along the Oregon coast have given rise to its irregular shore, with headlands of resistant basalt separating embayments formed by the

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⊡⊡ Fig. 1.3.4 The condominiums constructed on the remnant of terrace adjacent to the Jump-Off Joe Landslide, with slippage of the bluff having stressed the foundations so the windows were broken, evident in this 1981 photo.

more rapid retreat of the less resistant mudstones and siltstones. These embayments are the principal sites of beach sand accumulation, each constituting what is termed a littoral cell, basically a stretch of beach that for the most part is isolated by the large headlands that prevent the exchange of sand between adjacent cells: >Figure 1.3.5 is a typical example of a littoral cell. The individual cells contain different quantities of sand, apparent in the widths of their beaches, with the quantities depending on the existence of sand sources such as rivers or sea cliff erosion, which differ from cell to cell. A few cells have virtually no modern-day sources, the sand in its beaches instead being relict, having reached the cell thousands of years ago. In a study of the mineral contents of the sands within the cells we found that the metamorphic rocks of the Klamath Mountains in southern-most Oregon and northern California had been a prime source of sand, even to the cells along the central to northern Oregon coast where the numerous headlands now make it impossible for the Klamath sand to reach those beaches (Clemens and Komar 1988). We concluded that the Klamath derived sand had been transported north­ ­ward during times of lowered sea levels in the Ice Ages, unhindered by the presence of headlands when the shores were seaward on what today is the continental shelf. With the melting of glaciers and the rise in sea level, those shelf sands were pushed landward within the migrating beaches, having become trapped between headlands in the littoral cells roughly five thousand years ago. With the beaches being contained within littoral cells bounded by large rocky headlands, they are in effect pocket

beaches, even though they may have long shore lengths of 10 to over 100 km. In general, during the summer months the waves arrive predominantly from the west to ­northwest, and this causes a southward displacement of the sand within the cells. In contrast, the waves of winter storms mainly arrive from the southwest, moving the sand back to the north. As a result, there tends to be a seasonal northsouth oscillation of beach sand within the littoral cells, but with the long-term net littoral drift effectively being zero. The existence of this net-zero longshore transport of beach sand was evident when jetties were constructed during the early twentieth century on the inlets to estuaries and bays. In contrast to jetty construction in Southern California and along most of the U.S. East Coast, which did block a net longshore transport of beach sand so that it accumulated to one side of the jetties while erosion occurred in the down-drift direction, the jetties along the Oregon coast generally resulted in sand accumulation both to their north and south sides, locally where the shoreline is partially sheltered from the waves by the jetties (Komar 1997). Thus, jetties constructed on the Oregon coast have not been a problem in terms of having induced erosion and property losses. However one exception was dramatic, where early in the twentieth century the community of Bayocean was lost to erosion, but that occurrence can be attributed to only a single jetty having been constructed, not the usual pair, with the single jetty having caused the beach sand to be swept through the inlet into Tillamook Bay, lost from the sand spit on which the community had been developed (Komar 1997).

Oregon

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⊡⊡ Fig. 1.3.5 The Lincoln City Littoral Cell, with Cascade Head to its north and Cape Foulweather forming its southern boundary. Siletz Spit is positioned midway along the cell, with the remainder of the beaches backed by sea cliffs, the erosion of which is now the main source of sand to the beach as seen in (>Fig. 1.3.4); sand carried down the Siletz River is deposited within the Bay, so that it represents a minor contributor to the beach.

While the beaches along the Oregon coast are predominantly backed by sea cliffs, a number of sand spits are found separating the ocean from the estuaries and bays, as seen in (>Fig. 1.3.5) for Siletz Spit. These spits individually point either north or south, and some in close proximity and within the same littoral cell point in opposite directions; therefore, the direction of spit extension does not provide evidence for the existence or direction of a net ­littoral drift. While in total the progressive loss of property along the Oregon coast resulting from sea cliff erosion has probably been greater, more dramatic has been the erosion of developments on the sand spits, beginning with Bayocean and more recently with the major erosion that has occurred on Siletz Spit (>Fig. 1.3.6), Alsea Spit, and Netarts Spit, with the latter having impacted a State Park. These variable quantities of sand within the littoral cells and the resulting widths of their beaches depend on the existence of modern-day sources as well as the volumes of relict sand. This determines the rates of backshore property erosion, which varies from cell to cell, governed by the capacity of the fronting beach to buffer those properties from wave attack. The Oregon coast is one of the world’s most dynamic environments, with the extremes of its waves and tides accounting for occurrences of erosion like that seen in (>Fig. 1.3.6). These extremes have a direct connection with the Earth’s evolving climate, including the intensification of storms and the waves they generate, which may be due to global warming, and the periodic occurrences of

major El Niños. The Oregon coast is noted for the severity of its winter storms, which typically generate waves having deep-water significant wave heights (the average of the highest one-third of the waves) greater than 10 m, and with the significant wave heights during the most severe storms having reached 15 m, at which time the highest individual waves would have been about 25 m, the height of a 10-storey building. Daily measurements of waves off the Oregon coast have been collected by buoys since the mid-1970s. Of concern, those measurements demonstrate that the wave heights on an average have been progressively increasing, with the significant wave heights of the strongest winter storms back in 1975 having been about 9 m, having increased from 12 to 15 m in recent years (Allan and Komar 2006). While our analyses were based on wavebuoy measurements, data for the storm intensities in terms of wind speeds and atmospheric pressures extend further into the past, and demonstrate that the increases in wave heights likely began at least as early as the mid-twentieth century. Although the exact cause is uncertain, the increases in storm intensities and wave heights may be associated with global warming, though it has also been suggested that particulate pollution in the atmosphere may be important, drifting across the Pacific from China. The increases in deep-water wave heights and periods measured by the buoys off the Oregon coast have produced parallel increases in the processes active on its beaches, in particular the sizes of the breaking waves and elevations

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⊡⊡ Fig. 1.3.6 The erosion of Siletz Spit during the winter of 1972–73, with one home being lost and others on promontories of riprap placed to protect them. The erosion was caused by the coincidence of high storm waves and elevated tides, enhanced by seaward flowing rip currents that locally cut embayments into the beach.

reached by the swash runup of the waves when they reach the shore (Komar and Allan 2002). The runup elevations on beaches are particularly important in that they combine with the high tides to produce erosion of dunes and cliffs backing the beaches. Our analyses have shown that the progressive increase in wave heights has resulted in parallel increases in average swash runup levels at rates that are greater than the global rise in sea level, a factor that undoubtedly has played an important role in the increased property erosion experienced along the Oregon coast. Oregon’s tides are classified as mixed: there usually are two highs and two lows each day, but with the highs reaching different levels. With an average range of about 2 m and a maximum spring tide range of 4 m, they are further classified as mesotidal. These are the predicted astronomical

tides based on the forces of attraction of the moon and sun on the ocean’s water. The actual measured tides on the Or­­ egon coast can differ significantly from those ­predictions, with the difference primarily being of interest when the measured tide is substantially higher than predicted, since such occurrences can result in beach and property erosion. One cause of elevated measured tides is the occurrence of a storm surge, created by the onshore-directed winds and low atmospheric pressures of major storms. Measurements of surges on the Oregon coast show that they elevate tides of the order of 1.0–1.5 m, and although they are much smaller than storm surges produced by hurricanes along the U.S. East Coast, due to the low slopes of Oregon’s beaches its storm surges shift the mean-water shoreline landward by some 25–40 m, increasing the impact of the storm-wave runup on shore-front properties. The most significant climate event in terms of its ­erosion impacts along the Oregon coast has been the ­occurrence of a major El Niño, like those during the winters of 1982–83 and 1997–98 (Komar et al. 2000; Allan and Komar 2006). Particularly noteworthy is that an El Niño significantly elevates the measured tides, on average by about 0.30 m but achieving a maximum difference of about 0.60 m between the measured and predicted tides. This is documented by assessments of the monthly averages of the increased tidal elevations, with the maximum occurring during December and January, corresponding to the months that tend to have the greatest numbers and intensities of storms. The increased water elevation itself can be accounted for in part by the thermal expansion of the coastal water, which even in normal years is warmer during the winter due to the presence of cold water in the summer caused by upwelling, the water achieving still higher temperatures and levels in an El Niño winter. Another component results from the northward-flowing coastal currents, with their deflection to the right by the Coriolis force acting to pile water up along the shore; again, which tends to be stronger during an El Niño, resulting in elevated monthly-averaged water levels. Of im­­ portance to the resulting erosion of the Oregon coast, this  increase in water levels spans the entire winter, in effect representing a sudden increase in mean sea level, even though it later returns to normal when the El Niño ends. This rise in the monthly-averaged water levels during an El Niño elevates the water at all stages of the tides, so for several months there is an appreciably enhanced ­probability that the runup of storm waves on beaches will impact shore-front properties. The occurrence of a major El Niño also tends to, on average result in a degree of increase in the heights of winter storm waves that reach the Oregon coast, but far more

Oregon

important is that the waves arrive more from the southsouthwest than during normal years. This is because the tracks of the El Niño storms are shifted more to the south as they cross the North Pacific, passing over the shores of California rather than Oregon and Washington as they do during normal winters. The result is that unusually large quantities of the beach sand within the littoral cells are transported northward during an El Niño winter, resulting in ‘hot spot’ erosion at the south ends of the cells, to the north of the headlands, as depicted in >Fig. 1.3.7. Furthermore, the enhanced northward longshore sediment transport also tends to deflect the inlets to bays and estuaries, forming ‘hot spot’ erosion sites to the north of those inlets (or where present, north of jetties that act as mini-headlands). Following the El Niño winter the sand slowly returns to the south within the littoral cells, eventually reestablishing the long-term equilibrium sand volumes, although this return may take several years during which the erosion within the hot-spot zones persists. Erosion during the major El Niños of 1982–83 and 1997–98 extended along the entire U.S. West Coast, with that in northern California, Oregon and Washington ­having primarily occurred in hot-spot zones as depicted in >Fig. 1.3.7. Furthermore, along the Oregon coast the ­1997–98 El Niño erosion was significantly expanded ­during the following winter of 1998–1999 when several storms generated waves that exceeded what then had been ⊡⊡ Fig. 1.3.7 The contrasting reversing longshore sand movements within Oregon’s littoral cells during a Normal Year versus a major El Niño when strong storms from the southwest result in an enhanced northward transport and the development of ‘hot spot’ erosion sites.

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projected to be the 100-year extreme event (later to be reanalysed and increased to a 16-m significant wave height). The combined impacts of those two winters were extensive, to both private properties and State parks, in what we characterised as having been a ‘one-two punch.’ >Figure 1.3.8 shows the erosion at Cape Lookout State Park on Netarts Spit, an example of hot-spot erosion in that the Park is located at the south end of the littoral cell, to the north of Cape Lookout. The erosion began during the 1982–83 El Niño (Komar 1997), at which time the high tree-covered dunes were eroded away, so the public bathrooms within the campground landward of the dunes were threatened. Although they were ­temporarily protected by the ­placement of riprap, it proved insufficient during the 1997–98 El Niño and major storms that swept across the campground during the following winter, damaging the bathrooms to the extent they had to be torn down. The erosion of sea cliffs and foredunes along the Oregon coast has led to the proliferation of riprap revetments, with an example being those on Siletz Spit seen in >Fig. 1.3.6. The use of a ‘hard’ structure was undesirable in Cape Lookout State Park following the El Niño erosion, so it was decided to follow a ‘design with nature’ approach by constructing an artificial dune containing a core of sand-filled geotextile bags and planted with native dune grass, fronted by a ­constructed cobble berm (‘dynamic revetment’), in essence a nourished cobble beach as found naturally along

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⊡⊡ Fig. 1.3.8 The 1982–83 El Niño erosion at Cape Lookout State Park, with the loss of the high tree-covered dunes and vertical supports of a failed log sea wall seen in the background. The bathroom was lost during the 1997–98 El Niño.

the Oregon coast, thus acting to dissipate the swash of the waves before they impact the rebuilt dunes. The resulting ‘structures’ (>Fig. 1.3.9) have the appearance of their natural counterparts along the coast, and thus far have provided protection to the State Park (Allan and Komar 2004). The use of ‘hard’ structures – riprap revetments and sea walls – along the Oregon coast has had adverse impacts, especially where they have been used to prevent the erosion of sea cliffs. In some littoral cells the cliff erosion is the primary source of sand to the beach, in locations where the cliffs are composed of uplifted Pleistocene beach and dune sand; for example, the erosion at Gleneden Beach (>Fig. 1.3.3) within the Lincoln City Littoral Cell (>Fig. 1.3.5) south of Siletz Spit. There has been a considerable proliferation of structures within that cell, to the extent that the cliff-erosion source of sand to the beach has largely been cut off. Based on the previous rates of cliff retreat within that cell, it was estimated that the quantities of sand added to and building up the beach effectively balanced the relative rise in the sea level at that site (Shih and Komar 1994). With the loss of that sand source due to the installation of shore-protection structures, this equilibrium no longer exists, with the expected prolonged detriment of the beaches within that cell. Impressive accumulations of dune sands are found on the Oregon coast estimated to be present along about 45% of the coast, either in the form of foredunes backing the

beaches or contained within the massive Oregon Dunes Recreation Area that extends from Florence, south to Coos Bay. People have had a major role in altering the vegetationcover of the dunes, which in turn has affected their morphologies, with mixed consequences. When EuroAmericans first settled the Clatsop Plains south of the Columbia River in the nineteenth century, the extensive dune fields were covered by dense grasses. Those native grasses could be eaten by livestock, and overgrazing quickly reactivated the dunes so that by the 1930s some three thousand acres of sand had become mobile. In 1934 this area was planted with European beach grass, which livestock will not generally eat. Its introduction has had unforeseen consequences as it rapidly spread along the coast. On the positive side, it captured sand blowing landward from the beaches, building up substantially higher foredunes than had previously existed with the native grasses, providing a greater degree of protection from erosion and flooding to backshore properties. However, this growth of the foredunes has had negative environmental consequences, particularly to the nesting of Snowy Plover that need areas of open sand; in recent years extensive efforts have been undertaken to locally remove this invasive dune grass, to provide nesting habitat for this endangered shore bird. The arrival of the European beach grass has also had negative consequences for the Oregon Dunes Recreation Area. A century ago those dunes existed as an unvegetated

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⊡⊡ Fig. 1.3.9 The reconstructed dunes with a core of sand-filled geotextile bags, and cobble berm at Cape Lookout State Park, used as an alternative form of shore protection (Allan and Komar 2004).

sand surface extending from the ocean shore to the ­precip­itation ridge of the dunes at their landward edge (Cooper 1958). Sand was free to blow inland from the beach to support the continued growth of the high dunes. However, since the arrival of the European beach grass in the 1930s, large foredunes have grown immediately landward from the beach, where they now capture the sand blowing inland, preventing it from reaching the inland dunes. The impact of that loss was first noted in the area immediately landward from the foredunes, where the ground level was lowered to the water table, permitting the growth of shrubs and other vegetation where the high dunes had previously existed. The aerial extent of the active dunes has substantially decreased, and there is concern regarding their long-term preservation. An attempt was made to remove the beach grass and foredunes along a portion of the Oregon Dunes, but that experiment has not been followed up by a larger-scale implementation of this potential solution.

References Allan JC, Komar PD (2004) Environmentally compatible cobble berm and artificial dune for shore protection. Shore & Beach 72:9–18

Allan JC, Komar PD (2006) Climate controls on US West Coast erosion processes. J Coast Res 22:511–529 Atwater BF (1987) Evidence for great Holocene earthquakes along the outer coast of Washington state. Science 236:942–944 Clemens KE, Komar PD (1988) Oregon beach-sand compositions produced by the mixing of sediments from multiple sources under a transgressing sea. J Sediment Petrol 56:15–22 Cooper WS (1958) Coast dunes of Oregon and Washington. Geol Soc Am Memoir 72:169 Komar PD (1997) The Pacific Northwest coast: living with the shores of Oregon and Washington. Duke University Press, Durham, NC, p 195 Komar PD (2004) Oregon’s coastal cliffs: processes and erosion impacts. In: Hampton MA, Griggs GB (eds) Formation, evolution and stability of coastal cliffs – Status and trends. US Geological Survey Professional Paper 1693:65–80, Washington, DC Komar PD, Allan JC (2002) Nearshore-process climates related to their potential for causing beach and property erosion. Shore & Beach 70:31–40 Komar PD, Shih SM (1993) Cliff erosion along the Oregon coast: a ­tectonic – Sea level imprint plus local controls by beach processes. J Coast Res 9:747–765 Komar PD, Allan J, Dias-Mendez GM, Marra JJ, Ruggiero P (2000) El Niño and La Niña: erosion processes and impacts. In: Proceedi­ngs of 27th International Coastal Engineering Conference, Sydney, New South Wales, Australia, pp2414–2427 Satake K, Shimazaki K, Tsuji Y, Ueda K (1996) Time and size of a giant earthquake in Cascadia inferred from Japanese tsunami records of January 1700. Nature 379:246–249 Shih SM, Komar PD (1994) Sediments, beach morphology and sea cliff erosion within an Oregon coast littoral cell. J Coast Res 10:144–157

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1.4 California

1.  Introduction The California coast is about 2,900 km long, which includes the San Francisco Bay and the offshore islands. Some 70% of the coastline is rocky or has cliffs,  although pocket and fringing beaches are found commonly below the cliffs. The rest comprises either sandy fringing beaches backed by dunes, as south of San Luis Obispo, barrier beaches fronting lagoons and wetlands, as near Eureka, or the wetlands around San Francisco Bay. As on the Oregon coast there are numerous islands and stacks in the coastal waters of northern California that are scattered residuals of hard rock left by the recession of cliffs cut into softer sediment. About 85% of the coast is actively eroding. Although long stretches of exposed west-facing shores remain largely natural, extensive stretches around San Francisco, Los Angeles, and San Diego have been much altered by development. Of the original 80,000 ha of wetlands (excluding San Francisco Bay), 52% have been destroyed by dredging and filling, and a further 40% have been moderately or severely damaged. The coast embraces five geomorphic provinces, the Klamath Mountains, Northern Coast Ranges, Southern Coast Ranges, Transverse Ranges, and Peninsular Ranges. These provinces are structural units whose geology and relief reflect post-Palaeozoic interaction between the ­westward-moving North American lithospheric plate and subducting plates farther west. California’s structural framework began to emerge in late Jurassic and Cre­taceous times as the Nevadan Orogeny caused the ancestral Kalmath Mountains, Sierra Nevada, and Peninsular Ranges to rise from shallow seas and an island arc along the western margins of the North American plate, accompanied by westward thrusting, volcanism, and batholith emplacement. Subduction of marine sedimentary and ­volcanic rocks in the deep eugeosynclinal trench along this plate margin, followed by uplift, produced the Franciscan Formation, a heterogeneous jumble of readily erodible greywackes, shales, and metamorphic rocks found throughout the Coastal Ranges. During late Cretaceous and Cainozoic times, great

thicknesses of clastic sediment and submarine volca­­nic rocks accumulated in subsiding basins along the ­continental margin and strike-slip faulting became significant. Orogenic activity culminated in the uplift of coastal mountains from Miocene to mid-Pleistocene times. Coastal lowlands were restricted to subsiding intermontane troughs (Ventura Basin), to pull-apart structures (Santa Maria Basin), or to the seaward ends of strike valleys (Eureka Basin). During Neogene times, the North American plate approached and then overran the East Pacific Rise, a complex spreading centre in the Pacific plate. As that portion of California west of the San Andreas Fault transferred to the western limbs of this spreading centre, the Peninsular Ranges mini-plate (including Baja California) accelerated northwest from 2 to 5 cm/a year causing the Transverse Ranges to rotate into their present east–west alignment. Reflecting the influence of en echelon transform structures in the subducted East Pacific Rise, extensive strike-slip faulting with locally intense thrust faulting came to dominate coastal zone structures and landforms. Tectonism has persisted throughout Quaternary times, as shown by deformed and faulted marine terraces and the  subsiding coastal basins. A spectacular flight of 13 marine terraces are found on the Palos Verdes Peninsula, the highest at 380 m above sea level being equated with the nearby mid-Pleistocene deposits (Woodring et al. 1946; Wehmiller et al. 1977). A similar terrace sequence are found on San Clemente Island to the south, where the straight cliff along the NE coast is probably a fault scarp. On San Nicolas Island farther west, the 240-m terrace   dates back to the mid-Pleistocene age, while a 120-m terrace is dated as far back as 190,000 years. Fragments of high terraces are indeed found throughout the coast-from the deformed sequence near Eureka, along the seaward slopes of the Coast Ranges, to the extensive mesas around San Diego where the 90–150-m Linda Vista terrace and its fossil barrier beaches are considered to be a million years old. However, it is the last (Sangamon) interglacial terraces that provide the most continuous record of Quaternary tectonism.

Edited version of a chapter published by Anthony R. Orme in The World’s Coastline (1985: 27–36). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Terraces at 6 m and -14 m near San Diego have been dated at 120,000 and 80,000 years  (Kern 1977); a 23–45-m terrace at Palos Verdes has been dated at 85,000 years; 30-m terraces along the Malibu coast have ages of 130,000 and 105,000 years; while a Holocene terrace forms part of an impressive sequence west of Ventura, where uplift of 10 m/1,000 ayear has characterised the past 0.6 million years (Yeats 1978; see Orme 1980 for comprehensive bibliography). Terrace uplift and deformation is also impressive between Santa Cruz and San Francisco. Evidence for coastal submergence (by eustatic, isostatic or tectonic movements) is also compelling. The Ventura Basin, for example, subsided tectonically at rates up to 9.5 m/1,000 year from 2.0 to 0.6 million years ago, allowing a 7-km thickness of Plio-Pleistocene sediment to accumulate beneath the Oxnard Plain. More recently the Flandrian marine transgression entered the Golden Gate about 10,000 years ago and, rising at a rate of 2 cm/a year, flooded laterally across San Francisco Bay as rapidly as 30  m/a year until 8,000 years ago (Atwater et al. 1977) (>Fig. 1.4.3). The rate of sea level rise then declined and for the past 6,000 years has averaged 0.1–0.2 cm/a year. During this time, however, Holocene salt marsh deposits within the Bay have undergone 5 m of tectonic and possibly ­isostatic subsidence, a rate of 0.8 mm/a year superimposed  onto the Flandrian transgression eustatic effects. Subsidence is even more pronounced along the San Andreas Fault zone through the Bolinas Lagoon (Berquist 1978). Over the past century, tide gauge data show that the sea level has been rising at an average rate of 1.5 mm/a year along much of the coast (Hicks and Crosby 1974). Thus, despite ample evidence for Quaternary uplift and deformation, recent coastal changes must be viewed against the Flandrian transgression and a continuing sea level rise. Tides, winds, waves, currents, tsunamis, mass wasting, sediment discharge, and human activity all influence the patterns of coastal erosion and deposition. Mean spring tide ranges are between 1.5 and 1.8 m along the coast maximum tides attaining 2.5 m. Tidal currents are strong through narrow straits. At the Golden Gate, the flood inflowing tide reaches 1.7 m/s and the ebb 2.3 m/s, augmented by the head of stream discharge through San Francisco Bay. At the mouth of Tomales Bay, the ebb current clashes with incoming swells to form dangerous 3-m sneaker waves over the bar. Prevailing winds reach the coast from the northwest, producing waves and determining dune orientation. North of Point Conception 30–50% of ocean swell arrives from the northwest and most ­others from WNW. South of Point Conception, changing coastal orientation, strong refraction, and offshore islands

cause 70% of swells to pass up the Santa Barbara Channel from due west, while 80% of swells approach Los Angeles from the WSW. Here also, southerly swells set up by late summer hurricanes off western Mexico, by Southern Hemisphere winter storms, and by local winter depressions passing along more southerly tracks may cause erosion on south-facing beaches. Wave heights at Ventura average 1 m but range from 0.3–7.0 m (Orme 1982). The predominant northwest swells set up strong longshore currents, up to 2 m/s, and longshore drift from north to south, although northward drift may be favoured by shore configuration and reversing currents. The effect of the 1,000 km wide cold California Current (mean velocity 0.1–0.3 m/s; net discharge 11 million m3/s) is more climatic and ecological than geomorphic, but fog affects weathering and soil moisture on coastal slopes and the ocean current and its inshore counter current can transport fine sediment flushed out of coastal rivers. Tsunamis occasionally affect the coast, an example being the 1964 Alaska earthquake being followed by a series of up to 7 m waves on the coast at Crescent City. Stream sediment discharge, so important for nourishing California’s beaches, reflects the availability of rock waste, the precipitation and runoff patterns, and the extent to which fluvial transport has been blocked by dams. Along the northern California coast, erodible rocks, steep slopes, and high precipitation and runoff combine to ­produce frequent landslides and high erosion rates. The Klamath, Eel and Russian Rivers together account for 77% of all fine-grained sediment discharged to the sea north of Point Conception. Sediment yielded by the Sacramento-San Joaquin river system are mostly deposited before reaching the ocean. Farther south, Franciscan rocks and Cainozoic sediment are similarly erodible and prone to mass movement, but whether this debris reaches the coast depends on the winter precipitation regime, notably on the frequency of high magnitude storm flows. For example, most of the 9 million tonnes of sand delivered to the shore by the Santa Clara River between 1933 and 1938 arrived in 6 days of floods in 1938. Further, sediment discharged by this river in the 1969 floods was 47.6 million tonnes, compared with 1 or 2 million tonnes in relatively dry years. The coastal sediment budget of southern California thus sees years of plenty (notably 1969, 1978, 1980 and 1983) and years of famine (the 30 years preceding the 1969 floods). Mass movement from sea cliffs and coastal slopes is also an important sediment source, notably where fractured or poorly consolidated rocks are found, but this source has been both aggravated and curtailed by road and other construction

California

work, for example along the Malibu coast west of Los Angeles and on the Big Sur coast. In many places, for one reason or another, Pleistocene landslides and slumps have been reactivated in recent times, often through changes in groundwater hydrology. Offshore California is represented north of Point Conception by a shelf and slope, and southward by a continental borderland. The relatively narrow northern shelf widens to 50 km off the Golden Gate but is dissected by several submarine canyons, notably Delgada, Noyo, and Bodega canyons off northern California and Sur, Lucia, and Arguello canyons farther south. The largest of all is the Monterey Canyon, which heads within 0.8 km of the shore but dissects the shelf and slope to a depth of over 3,500 m. A large lunate sand bar composed of terrestrial sediment lies off the Golden Gate. The continental borderland off southern California comprises a series of fault blocks and troughs, with some closed basins and rugged emergent islands. These structures trend east-west off the Transverse Ranges and northwest-southeast farther south, and their origins must be related to the Cainozoic behaviour of the adjacent land areas (Emery 1960). At least 32 distinct submarine canyons dissect southern California’s offshore area.

2.  The Coastline of Northern California South from the Oregon border the coast consists of low cliffs and bluffs fringed by grey sandy beaches, with a few ⊡⊡ Fig. 1.4.1 Stacks and bluffs in the bay south of Trinidad Head. (Courtesy Geostudies.)

1.4

islands offshore. Lake Earl is a lagoon north of Point St George, where the bluffs steepen and are undercut by low cliffs. The northern California cliffs and bluffs are cut in pulverised eugeosynclinal Franciscan rocks consisting of greywackes and shales, and though they are forested they are prone to frequent mass movement and rapid erosion. These rocks strike NW-SE and have been eroded into narrow valleys that follow the strike and open into bays separated by elongated ridges that end in headlands. On this exposed coast winter storms generate waves 4–7 m high  at Point St George where the coast turns southeast, and Cres­ cent City stands on a low plateau edged by cliffs, with a series of headlands and bays. The tsunami of 27 March 1964 sent four waves, the highest over 6 m, in through the town, demolishing 29 city blocks. Sandy barriers enclose Freshwater Lagoon, Stone Lagoon and Big Lagoon, separated by spurs of high ground. The lagoons have rushy marshes and are backed by steep slopes. At Trinidad the coast is steep and partly cliffed, and to the south the coast is fringed by numerous islets and stacks, consisting of hard chert, greenstone and sandstone eroded out of a matrix of soft mudstone in the Franciscan Mélange (>Fig.  1.4.1). One of these harder components protrudes from a cliff of slumping clay at Luffenholtz, and will in due course become a headland, then an island. The steep coast ends as the Little River flows out at the northern end of a long sandy beach. Humboldt Bay is a large coastal lagoon behind a dunecapped sandy barrier, with a central outlet bordered by jetties. The shores are partly marshy (> Fig.  1.4.2), and

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⊡⊡ Fig. 1.4.2 Flat-topped cliffs and stacks indicate the coastal terrace at Ten Mile River. (Courtesy Geostudies.)

shoreline erosion is extensive, possibly as a result of the stabilisation and dredging of the entrance to allow ships access to Eureka. The River Eel carries large quantities of sand and gravel down to an estuary largely filled with shoals, but it has not yet built a delta, and the river mouth opens between paired spits. A long steep and cliffy coast, trenched by deep valleys incised into the coast ranges, extends south, past Cape Mendocino to Punta Gorda, then southeast to Point Delgarda and Cape Vizcaino. There are ­sloping cliffs and sandy beaches, as at Union Landing. Headlands end in lines of stacks near Westport and there is a wide coastal terrace which extends on to flat-topped islands at Brutal Point. The even-crested cliffs end at Ten Mile River (> Fig. 1.4.2) and a long sandy beach is backed by dunes, some drifting inland. At Fort Bragg the cliffs return, fronting a coastal terrace, and there are again many rocky stacks. Glass Beach is unusual in that it is derived from broken glass, porcelain and other garbage dumped from the cliffs, a procedure that was halted in 1970. To the south, on the coastal terrace, is Mendocino, with cliffs in sandstone and clay and reefs with kelp at the mouth of Big River. There are beaches of grey sand and pebbles at the mouth of Little River, and the coastal terrace continues southward, backed by a steep rising slope. It was cut when sea level stood higher relative to the land, and rocky knolls that rise from the terrace were formerly stacks. Navarro River is one of several rivers that have incised deep steep-sided valleys across the coastal terrace. It has a small estuary, encumbered by sand spits with some lumber. At Alder Creek the cliffs pass

inland as bluffs behind a sandy beach and dunes, which extend to the mouth of Hathaway Creek at Point Arena. The sloping cliffs at and south of Point Arena are receding evenly, cut in soft shale over sandstone and rising to a broad coastal terrace. In places the basal sandstone forms shore platforms, usually coinciding with bedding planes, while interspersed outcrops of hard grey gnarled rock protrude as headlands and are eventually isolated as reefs and stacks. Inland, the Mendocino Range runs parallel to the coast, and there are segments where the coastal terrace fades out to a steep coast, as at Saunders Landing. Rivers deliver gravelly loads to the shore, and beaches are mixtures of sand and pebbles. South from Stewarts Point the cliffs are cut in yellow sandstones which show honeycomb weathering, and shore platforms are well developed near Walsh Landing. Fort Ross stands on a coastal terrace that slopes gently seaward, incised by short narrow steep-sided valleys and ending in cliffs and coves (> Fig. 1.4.3). The coastal terrace is interrupted by a steep sector, the Sonoma Coast, but resumes to be crossed by the Russian River, which descends from a deep gorge cut through the recently uplifted Coast Range. The estuary is blocked by a barrier of grey sand (> Fig. 1.4.4), through which the river generally maintains an outlet that is enlarged during flood discharge. Near Peaked Hill, former stacks rise abruptly from the coastal terrace, which is backed by slopes marking a degraded cliff. Shell Beach and Duncans Beach are wide in front of cliffs, but to the south the beach narrows and rapid cliff recession threatens houses at Sorano del Mar.

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1.4

⊡⊡ Fig. 1.4.3 Coastal terrace and cliffs at Fort Ross. (Courtesy Geostudies.)

⊡⊡ Fig. 1.4.4 Sand barrier at the mouth of Russian River. (Courtesy Geostudies.)

From Salmon Creek the sandy beach resumes, and becomes a barrier backed by dunes and curving out to Bodega Head. On the landward side is Bodega Harbour, with a narrow outlet at the western end of a curving sand spit, and the coast then runs close to the line of the San Andreas Fault along Bodega Bay. Dillon Beach is wide and sandy, receiving ocean swell, but Tomales Point, offshore, marks the beginning of the long narrow Tomales Bay, where wave energy is much reduced. There are successive

headlands and bays along the coastline that runs southeast, past the salt marshes in the Chileno River estuary (> Fig. 1.4.5) and along the seaward slopes of the Balinas Ridge. Bluffs pass into cliffs on the more exposed promontories, south of Marconi, and at the head of Tomales Bay there is a marshy delta at the mouth of Olemo Creek. The inner shore of the Point Reyes Peninsula is fairly straight, close to the San Andreas Fault, and consists of Mesozoic granite. Lateral movement along this fault is

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⊡⊡ Fig. 1.4.5 Salt marsh in Chileno River estuary, Tomales Bay. (Courtesy Geostudies.)

causing the Point Reyes Peninsula to move northwest along the coast at an average rate of about 5 cm/a year, much of the movement occurring during earthquakes, as in 1906 when there was sudden displacement of up to 6 m. The outer coast at Tomales Point is exposed to oceanic waves, and cliffs cut in red sandstone underlain by hard grey Sierran granite extend south past McLures Beach. The high energy Pacific swell generates major rip currents here. Much of the west coast of the Point Reyes Peninsula has a wide sandy beach backed by dunes, forming a barrier that impounds Abbotts Lagoon in a gently incised valley. At the southern end, Point Reyes is a steep-sided ridge running west to east, and on the south coast is the broad curve of Drakes Bay, with cliffs cut in soft Pliocene sediment (>Fig. 1.4.6) and beaches that become barriers and spits in front of lakes (Drakes Estero and the adjoining Estero de Limantour) impounded in rias. Drakes Head is a cliffed promontory behind the barrier spits, and an outlet from the esteros has an ebb tide delta that refracts incoming waves. The coast steepens on the flanks of forested ranges down to Bolinas Point and Duxbury Point. U.S. Coast Survey records indicate that these points have been receding at about 0.5 m/a year. The San Andreas Fault runs southeast from Tomales Bay along a valley that ends in Bolinas Lagoon, which is almost enclosed by a sandy spit. The steep forested coast then continues to Point Bonita, on the northern side of Golden Gate, the strait at the mouth of San Francisco Bay.

3.  San Francisco Bay San Francisco Bay occupies a late Pliocene to mid-­Pleistocene structural trough within the Coast Ranges. The trough has a Franciscan basement bounded by ­northwest-trending strike-slip faults and is largely filled with Neogene and Quaternary marine and fluvial sediment. Continuing tectonism apart, this trough has been ­further deepened by fluvial erosion and invaded several times by Quaternary transgressions (Atwater et al. 1977). The bay is connected to the Central Valley through the Carquinez Strait, with a bedrock channel incised by the ancestral Sacramento River to 60 m below sea level, and to the Pacific Ocean through the Golden Gate, with a channel that is scoured to 104 m below sea level. San Francisco Bay is a natural sump for an 80,000 sq. km drainage area comprising the Sacramento-San Joaquin system. Some 200 years ago the bay waters covered 1,800 sq. km, but they have been reduced to 1,100 sq. km, and 70% of the bay is shallower than 4 m. This reduction is due in part to massive fluvial sedimentation during and after a period of hydraulic gold mining in the Sierra Nevada foothills in the latter half of the nineteenth century, and in part to extensive coastal land reclamation. Gilbert (1917) estimated that 1,816 million m3 of debris were eroded in the foothills by mining and natural processes between 1850 and 1914, of which 50% was deposited in the valley downstream, 48% reached the bay and only 2% reached the ocean. He also estimated that it would take 50 years (until 1964) for the upstream debris to reach the bay, after which the yield

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⊡⊡ Fig. 1.4.6 Cliffs in Drakes Bay, Point Reyes Peninsula. (Courtesy Geostudies.)

would diminish to about 6 million m3/a year, a figure that accords well with recent estimates. Wave action within San Francisco Bay is limited by the short fetch, and breaker heights average at 0.3 m. Mean spring tide range is about 1.8 m, and tidal currents rarely exceed 1 m/s within the bay. The salinity averages 27–29‰, with dilution by inflow of about 570 m3/s from the Sacramento-San Joaquin River. The relatively low-energy shores have salt marshes around mean high tide level dominated by Spartina leiantha, Distichlis spicata, and Salicornia ambigua zones, upper tidal flats between mean high and mean low water and lower tidal flats below mean low water (Pestrong 1972). About 200 sq. km of marshes and mudflats are exposed at mean tide level.

4.  The Coastline of Southern California It is convenient here to consider Southern California as the coast, south from San Francisco. The Southern Coast Ranges extend 400 km southeast from the Golden Gate to Point Arguello and are dominated by massive deformation and faulting, strike valleys, and elongated ridges. Between the San Andreas and Nacimiento fault zones, metamorphic and granitic rocks form the 60-km-wide Salinian basement. These rocks reach the coast in rugged slide-prone cliffs south of Pacifica. Pillar Point is a southward projection bordering Half  Moon Bay, which has been shaped by ocean swell refracted around the point and has a beach showing lateral

gradation in grain size, coarsening and steepening as exposure to wave action increases southward (Bascom 1951). A generally steep coast with valleys incised into the ranges extends past Pigeon Point and Point Ano Nuevo, where the coast curves to Santa Cruz in Monterey Bay. Only one of the natural arches cut in Santa Cruz mudstone is still standing; the others have collapsed to leave stacks offshore. Late Cretaceous and Tertiary sediment reach the coast north of Monterey Bay in deformed marine terraces and eroding sea cliffs (Bradley and Griggs 1976). The terraces have been rising at 1.6–2.6 cm/century and long-term sea cliff erosion rates range from zero in resistant granodiorites to 0.3 m a year or more in softer Tertiary  sediment, reaching 0.6 m/a year locally (Griggs and Johnson 1979). Extensive sand dunes around Monterey Bay are ex­­ plained in part by the abundance of sediment formerly brought down by the Pajaro and Salinas rivers. For 150 km south from Monterey the Santa Lucia Range dominates the coast, often reaching 1,000 m within 2 km of the shore. Where Franciscan rocks reappear west of the Nacimiento Fault, landsliding is a perennial problem. South of Ragged Point, the mountains trend inland, emerged marine terraces reappear, and broad structural basins are veneered with extensive dunes of late Pleistocene and Holocene age, notably behind the 25 km of sandy beach between Pismo Beach and Point Sal. Sand for these dunes comes mainly from the Arroyo Grande and from the Santa Maria and Santa Ynez rivers. Cliff erosion threatens homes and roads at Cayucos, Avila Beach, and Shell

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Beach, despite partial shelter from northwesterly swells. In Morro Bay salt marshes are protected by a long barrier spit. In the hinterland is a series of transgressive dune formations (Orme 1990). The Santa Ynez Mountains, which rise to 1,430 m extend 125 km east from Point Arguello to Ventura, and are a south-dipping homocline of mostly Tertiary marine and fluvial sediment. Quaternary uplift has produced much deformed marine terraces along the coast. Though protected by offshore islands, refracted westerly ocean swell have produced average erosion rates of 0.15 m/a year in the relatively soft Neogene sediment along the Santa Barbara coast. A small barrier-lagoon system survives at Carpinteria. The Oxnard Plain is a triangular lowland, formed mainly by deposition from the Santa Clara River and Calleguas Creek, where the Ventura structural trough plunges westward into the Santa Barbara Channel. The coastline once comprised a sandy barrier beach with low dunes backed by lagoons and marshes, breached by rivers during winter floods, but there has been much construction along the shore. Breakwaters and groynes have been built at Santa Barbara, Ventura and Port Hueneme, and there have been problems with interception of eastward longshore drifting of sand by the prevailing westerly swell. About 750,000 m3/year of sediment moves along this coast in one of several littoral cells on the coast of Southern California, each with sand supplied mainly from rivers passing alongshore until it is lost into the head of a submarine canyon (Emery 1960). To the south of the Santa Barbara Channel the Channel Islands (Anacapa, Santa Cruz, Santa Rosa, San Miguel) are underlain by a metamorphic or granitic basement and mantled with thick late Cretaceous and Tertiary rocks, notably Miocene marine shales and volcanic formations. The Santa Monica Mountains, which rise to 949 m, extend 50 km along the coast between Point Mugu and Santa Monica and are bounded to the south by the active Malibu thrust fault, the impact of which is seen in seismic activity, deformed terraces, fractured coastal rocks, and frequent mass movement. Much of the coast is typified by  unwise housing development on the backshore or beneath crumbling cliffs, creating inevitable problems during storms and southerly swells. Access is restricted to many beaches, as at Malibu. The Peninsular Ranges extend SSE for 200 km behind Los Angeles to the Mexican border. This mini-plate, ­comprising basement rocks, granitic plutons and post-­ batholithic sediment, has been diverging from the North American mainland at a mean rate of 6 cm/a year since it shifted on to the western limb of the East Pacific Rise some 4–5 million years ago.

Along the coast deformed marine terraces such as those on Palos Verdes (> Fig. 1.4.7), San Onofre Moun­ ­tain  and around San Diego, and active faults as in the Inglewood-Rose Canyon structural zone that outlines the coast between Newport Bay and San Diego, attest to continuing tectonic activity. The beach fringed Los Angeles coast is heavily urbanised, and Long Beach is a port and industrial urban complex where coastal land subsidence has followed oil extraction. Barrier and lagoon systems were once common along this coast, but most have been modified, and some obliterated, by reclamation. Ballona Lagoon, once an outlet for the Los Angeles River, is now the marine  and Newport Bay, until 1915 the outlet for the Santa Ana River, was heavily developed. Offshore is Santa Catalina Island, which has warped emerged marine terraces. Much of the coast south of Newport is cliffed, between valleys where rivers supply sand to the shore, sometimes through small barrier lagoons. This is another littoral cell in which beaches have been described as ‘rivers of sand,’ fluvial sediment drifting southward along the coast until it disappears into a submarine canyon off San Diego. Sand is also derived from receding cliffs, partly by marine erosion and slumping and partly by the effects of runoff. Erosion is a significant problem along much of this coast. Powerful ocean swell and storm waves arrive through windows between the offshore islands, coastal rock formations are often poorly consolidated and, owing to floodcontrol dams and channelization along most of the rivers, fluvial sediment supply has diminished in recent decades, depleting protective beaches. Cliff erosion is particularly rapid in front of Camp Pendleton, south of Oceanside harbour, near Encinitas and Del Mar, and at Sunset Cliffs, San Diego. Groundwater discharge weakens soft cliffs, causing exudations and triggering slumping. Inappropriate clifftop development has generally aggravated the problem, for example, through irrigation and lawn watering, increased impermeable surfaces, and poorly located culverts. At La Jolla Scripps Pier has been used as a transect for measuring fluctuations in the beach profile, notably the cut and fill cycles that accompany alternations of winter storm erosion and summer shoreward drifting. These cycles can result in beach lamination, with layers of fine and coarse sediment deposited during accretion with phases of calm-weather fine sedimentation and rough weather winnowing. Rocky shore outcrops show evidence of weathering by alternations of wetting and drying, producing platforms at the level of permanent saturation, as well as corrosion bysalt weathering. Boomer Beach, near La Jolla, is well known for its seasonal alternations. In winter storm waves scour away

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⊡⊡ Fig. 1.4.7 Palos Verdes Bay, showing raised terrace. (Courtesy Geostudies.)

sand exposing a rocky and bouldery shore, but in summer constructive waves sweep sand back in from the sea floor to bury these beneath a sandy beach. To the south San Diego has an urbanised coast, and the former San Diego River outlet has become an important naval and commercial harbour, bordered by Coronado Strand, a built-over spit. The urbanised coast extends to the Mexican border.

References Atwater BE, Hedel CW, Helley EJ (1977) Late quaternary depositional history, Holocene sea level changes, and vertical crustal movement, Southern San Francisco Bay, California. US Geological Survey Professional Paper 1014, Washington, DC Bascom W (1951) Relationship between sand size and beach face slope. Trans Am Geophys Union 32:866–874 Berquist JR (1978) Depositional history and fault related studies, Bolinas Lagoon, California. US Geol Surv Open-File Report 78–802, Washington, DC Bradley WC, Griggs GB (1976) Form, genesis, and deformation of central California wave-cut platforms. Bull Geol Soc Am 87:433–449 Emery KO (1960) The sea off Southern California. Wiley, New York Gilbert GK (1917) Hydraulic-mining debris in the Sierra Nevada. US Geol Surv Prof Paper 105, Washington, DC

Griggs GB, Johnson RE (1979) Coastline erosion, Santa Cruz County, California. Geology 32:67–76 Hicks SD, Crosby JE (1974) Trends and variability of yearly mean sea level 1893–1972. US National Oceanic and Atmospheric Administration, Technical Memorandum NOS 13, Washington, DC Kern JP (1977) Origin and history of upper Pleistocene marine terraces, San Diego, California. Bull Geol Soc Am 88:1553–1566 Orme AR (1980) Marine terraces and quaternary tectonism, northwest Baja California, Mexico. Phys Geogr 1:138–161 Orme AR (1982) Temporal variability of a summer shore zone. In: Thorn CE (ed) Space and time in geomorphology. George Allen & Unwin, London Orme AR (1990) The instability of Holocene coastal dunes: the case of the Morro Bay dunes, California. In: Nordstrom KF et al (ed) Coastal dunes: form and process. Wiley, New York, pp 315–336 Pestrong R (1972) San Francisco Bay tidelands, California. Geology 25:27–40 Wehmiller JE, Lajoie KR, Kvenvolden KA et al (1977) Correlation and chronology of Pacific coast marine terrace deposits of continental United States by fossil amino acid and ­stereochemistry-technique evaluation, relative ages, kinetic model ages, and geological implications. US Geological Survey Open-File Report, Washington, DC, pp 77–680 Woodring WP, Bramlette MN, Kew WSW (1946) Geology and Paleontology of Palos Verdes Hills, California. US Geological Survey Professional Paper 207, Washington, DC Yeats RS (1978) Neogene acceleration of subsidence rates in southern California. Geology 6:456–460

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1.5 Texas

Robert Morton

1.  Introduction The Texas coast bordering the western Gulf of Mexico (>Fig. 1.5.1) is about 590 km long and consists of barrier islands (80%) and mainland beaches (20%). Of the sandy barrier segments, about 45% are retreating, 35% are progradational features and 20% are aggradational features.

Retreating barrier and beach segments, such as the Rio Grande deltaic headland and adjacent South Padre Island (>Fig. 1.5.2) are narrow, thin and have low topographic profiles that allow repeated overwash by extreme storms. These relatively young landward-migrating barriers are located on the flanks of eroding Holocene deltaic headlands constructed by the Rio Grande and the combined

⊡⊡ Fig. 1.5.1 The Texas coast consists of mainland beaches and barrier islands with associated bays and lagoons.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.5.2 South Padre Island is a narrow, low, retreating barrier that is flooded and overwashed frequently by hurricanes.

⊡⊡ Fig. 1.5.3 Central Padre Island is an aggradational barrier with wide stable or accreting beaches and high, vegetated sand dunes.

Brazos-Colorado rivers. In contrast, the aggradational barrier segments with associated tidal flats (>Fig. 1.5.3) and progradational barriers with associated washover fans (>Figs. 1.5.4 and > 1.5.5) are wide, thick and have high dunes that effectively prevent storm surges from washing over the barrier islands. These Holocene barriers formed within the interdeltaic embayments where sand was supplied by longshore currents. The western Gulf of Mexico was one of the first coastal regions, where systematic research regarding

modern sediment and coastal processes was conducted. These efforts, which began in the late 1940s, were largely driven by the petroleum industry that was interested in developing models for predicting subsurface sedimentary facies and reservoir characteristics. Classic field-based geologic studies of Holocene barrier islands (Fisk 1959; LeBlanc and Hodgson 1959), sea level history (Curray 1960) and sediment distribution (Shepard et al. 1960) used Texas coastal features as a research laboratory.

Texas

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⊡⊡ Fig. 1.5.4 Long, shore-parallel beach ridges on San Jose Island are characteristic of progradational barrier islands.

⊡⊡ Fig. 1.5.5 A large inactive hurricane washover fan on San Jose Island. (Courtesy Texas General Land Office.)

2.  Geological Setting Like other barrier coasts of the world, the Texas coast owes its position and shape to sea level fluctuations associated with expansion and contraction of continental ice masses during the late Pleistocene. During the last eustatic cycle, most coastal plain rivers maintained their courses, incising deep valleys during the falling phase and backfilling, as sea level rose to its current position. Bays

oriented perpendicularly to the modern coastline represent unfilled remnants of former incised valleys. Only rivers with large drainage basins, such as the Rio Grande and combined Brazos-Colorado systems, carried enough sediment to aggrade their upper valleys and prograde deltas of moderate size into the Gulf of Mexico. Although the general Holocene sea level curves show a gradual reduction in the rate of rise and relative stability beginning about 5,000 years ago (Curray 1960), there is convincing

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Texas

evidence that sea level was about 1–1.5 m higher than present during the past few thousand years (Morton et al. 2000). The beaches and barriers of the Texas coast are arranged in an orderly pattern alongshore (Morton 1979). The Holocene Rio Grande and Brazos-Colorado fluvial-deltaic systems and a headland composed of Pleistocene fluvialdeltaic deposits control the coastline types and their positions. These three headlands formed large promontories in the Gulf of Mexico that focused wave energy and created three cells of littoral drift convergence. The headlands are composed primarily of stiff mud (>Fig. 1.5.6) and the nonbarrier beaches that front the headlands are rapidly eroding (Morton et al. 2004). The longest barrier-island segment formed within the southernmost littoral drift cell, which extends from the Rio Grande delta to the Brazos-Colorado ⊡⊡ Fig. 1.5.6 Shores of the Holocene Brazos-Colorado delta, such as this one at Sargent, Texas, typically are composed of stiff, organic-rich mud overlain by thin washover terrace deposits.

delta. Retreating barriers within this segment, including South Padre Island (>Fig. 1.5.2) and Matagorda Peninsula, are located on the flanks of the deltaic headlands (Morton 1979). The barriers are migrating landward as sand eroded from the Gulf coastline is transported across the barriers through the washover channels and into the adjacent lagoons. Transitional barrier segments separate these migrat­ ing barriers from the progradational and aggradational barriers, including Central Padre Island (>Fig. 1.5.3) and North Padre Island, Mustang Island, San Jose Island, and Matagorda Island, which formed in the center of the southern interdeltaic embayment. Within the middle littoral drift cell, which formed between the Brazos-Colorado deltaic headland and the Pleistocene headland, sand-rich progradational barrier islands Galveston Island and Bolivar Peninsula and a thin transitional barrier (Follets Island) were deposited. The progradational barriers have moderately high beach-ridge elevations (3–4 m) and small sand dunes. Cheniers consisting of marsh mud alternating with narrow beach ridges of shelly sand, accumulated against the Pleistocene headland within the easternmost littoral drift cell. This coastline segment, which is west of Sabine Pass, forms the western extension of the chenier plain of southwestern Louisiana. Texas barriers are entering a new phase of their evolution because of recent increases in the rate of relative sea level rise and decreases in sediment supply (Morton and McGowen 1980). Retreating barriers continue to migrate landward, while most of the other barriers are getting narrower. This is because of simultaneous erosion along both Gulf and lagoon shores. An exception to this trend is Central Padre Island, which remains an actively aggrading barrier. Sand released by updrift beach erosion is transported to this barrier segment by converging littoral currents. The sand supplied to the Central Padre Island beaches, dunes and backbarrier environments (tidal flats) by longshore and aeolian transport, is sufficient to maintain a stable landform even as relative sea level rises.

3.  Climate The Texas coastal climate, which ranges from subhumid in the east to semiarid in the south, is greatly influenced by the systematic westward decrease in precipitation and increase in evapotranspiration. The strong climatic gradient has a dramatic affect on biological assemblages, landforms and physical processes (Morton and McGowen 1980). The average annual rainfall for the eastern coast is 125 cm. As a result, nutrient-rich freshwater inflow to the

Texas

bays supports lush saltwater marshes and large oyster reefs, but the higher rainfall also minimises the influence of aeolian processes. In south Texas, where evapotran­ spiration generally exceeds precipitation, coastal waters commonly have abnormally high salinity and saltwater marshes represent only minor bay-margin environments; moreover, large active dune fields and tidal flats dominate the nearshore landscape.

4.  Wind and Waves The Texas coast is a microtidal, storm-dominated region that is influenced periodically by energetic meteorological events (Hayes 1967). Most of the year, wind blows from the southeast, but during the winter, strong northerly winds accompany the passage of cold fronts that can suddenly lower the temperature, rapidly change wind direction and greatly increase wave heights. These winter storms are high-frequency events that occur several times each year. Highest sustained wind speeds accompany major hurricanes such as Carla (1961) and Allen (1980). These infrequent extreme storms are so large that their organised wind patterns fill the northwestern Gulf of Mexico, matching the scale of the coastline arc. The counterclockwise wind circulation of hurricanes, drives strong nearshore currents and sand transport to the west and southwest in the Gulf of Mexico. Fair-weather waves in water depths of 10 m are normally about 60 cm high and have periods of 6 s. The broad shallow continental shelf bordering the Texas coast causes deep-water swell in the Gulf of Mexico to decompose into these low, short-period waves. Although precise measurements of waves breaking along the South Texas coast have not been made, field observations indicate that wave heights are slightly greater along Central Padre Island where forebeach slopes and offshore gradients are slightly steeper than in adjacent areas. The largest deep-water waves recorded in the Gulf of Mexico were 27 m high and had period of about 16 s. These hurricane-generated waves break far from the shoreline because of the broad shallow shelf. The storm waves reform and repeatedly break, creating a wide energetic surf zone that encompasses as many as nine bands of breakers.

5.  Tides Astronomical tides in the western Gulf are diurnal or mixed and typically have a range of about 0.5 m during a normal tidal cycle. Tide ranges are even lower in bays and

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lagoons because of the small number of inlets and great distances from most inlets to the bay margins. Windinduced changes in bay water levels are commonly more significant than those caused by the astronomical tidal cycle (Kraus 2007). Lowest water levels in the bays and lagoons typically are associated with strong northerly winds associated with winter storms that drive water out of the bays. In contrast, highest water levels are caused by hurricanes. During some hurricanes, tidal inlets cannot maintain water levels between the lagoons and the Gulf. As a result, large hydrostatic differentials can be created across the barriers, and powerful currents erode deep washover channels through the barrier islands (Morton 2002).

6.  Present Sea Level Tide gauge records in Texas show the same general variations in sea level that coincide with droughts and periods of abnormally high rainfall. They also show a relative rise in sea level averaging 3.4 and 6.5 mm/year at Port Isabel and Galveston respectively. Some of the relative rise in sea level is caused by compaction of the basin fill and subsidence of the land surface. The high rates of relative sea level rise measured along the southeastern Texas coast are attributed to subsidence induced by shallow ground-water withdrawal and deep hydrocarbon production (Morton et al. 2006).

7.  Tidal Inlets Both primary and secondary tidal inlets are present along the Texas coast (Kraus 2007). A single primary inlet is located in the southern sector of each bay system because of the predominant wind directions and southwesterly sediment transport. Primary inlets are large permanent channels that serve as the principal conduits between the Gulf and large bay systems. The primary inlets are wide, short and oriented perpendicular to the Gulf coastline. This configuration makes them highly efficient trans­ porters of water and sand. They are also relatively stable, having occupied the same general positions during construction of the adjacent barriers. The secondary inlets are narrow, long and slightly oblique to the Gulf coastline. These moderately efficient but unstable inlets, migrated in the direction of net longshore transport before they were modified. A third category of inlets includes the small, ephemeral features that are occasionally active during hurricanes. They are highly inefficient and typically close a few years after being reopened.

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8. Beach and Barrier Island Characteristics Most Texas beaches are composed of well-sorted, subangular to subrounded, fine to very fine sand regardless of whether they are eroding or accreting. Exceptions are the highly erosional deltaic headlands, where mobile beach sediment forms a thin veneer over stiff mud. Locally, these beach deposits contain high concentrations of gravel-size terrigenous and biogenic detritus (Morton and McGowen 1980). Broken and whole shells constitute as much as 50% of some erosional beach deposits. The shells are of brackish to normal marine fauna (Rangia cuneata, Crassostrea virginica, Mercenaria campechiensis) that are diagnostic species of estuarine and bay environments. The non-biogenic gravel is composed of rock fragments and caliche nodules also eroded from the underlying sediment and concentrated on the beach. The gravel fraction of shells and rock fragments represents a coarse lag remaining, after constant winnowing and continuous erosion of sand-poor relict sediment exposed on the shoreface and inner shelf. High concentrations of shell also occur on Central Padre Island. These shells are of Gulf species that were left as a wind-deflation lag on an accreting beach fed by converging longshore currents. Elevations of Texas back-beaches are only about 1 m because of the small tide range and low wave heights. Storm surges generated by major hurricanes, commonly exceed the elevations of backbeaches by 2–3 m. Average wave characteristics and offshore slope produce an upper shoreface about 600 m wide where three migratory breaker bars are commonly maintained by the interaction of waves and longshore currents. Although much of the littoral drift is transported by traction and suspension in this zone of breaking waves, substantial volumes of sand are stored and transported across the middle and lower shoreface. The shoreface extends seaward of the surf zone to depths of at least 10 m and distances of more than 2 km offshore. During major storms, the surf zone widens, wave heights increase, longshore currents accelerate and sediment transport significantly increases along the shoreface. The Rio Grande serves as the international boundary between the United States and Mexico. It drains a large area that is mostly desert or has low rainfall. Withdrawal of water along its course for irrigation and industrial purposes has so severely reduced the river’s flow that the mouth is commonly closed by a shoal. Beaches near the mouth of the river are eroding because wave energy is moderately high and there is a lack of sediment supply. South Padre Island is a retreating barrier that forms the northern flank of the submerging Rio Grande delta.

The southern tip of South Padre Island is densely developed and the site of high-rise construction. The development on South Padre Island is progressing to the north into a hazardous area, where the island is low and narrow, the beach is eroding, the dunes are discontinuous and broad washover channels extend entirely across the island (>Fig. 1.5.2). Some buildings have already been constructed in washover channels. The semiarid climate of South Texas promotes aeolian sand transport and the migration of low active dunes over the road along South Padre Island. During extreme storms the dunes are eroded and become part of the washover sediment deposited in the backbarrier flat or in Laguna Madre. Padre Island, north of the Mansfield Ship Channel is transitional between the narrow migrating barrier of South Padre Island and the wide aggradational barrier of central Padre Island. Beach erosion and dune migration characterise this segment of the coast because the vegetation is too sparse to stabilise the sand. The most common backbarrier subenvironment along Padre Island is a tidal flat. These broad sand flats are typically barren except for a thin algal mat. Plants cannot survive this harsh environment that are sometimes exposed and dry, and sometimes flooded with marine (lagoon) water that becomes hypersaline as it evaporates. Central Padre Island is an aggradational barrier with stable or accreting beaches that receive sand delivered by longshore currents and aeolian transport. It is in the zone of convergence between currents flowing northward from the Rio Grande delta and southwestward from the central Texas coast. This barrier segment is characterised by high, continuous and densely vegetated dunes with elevations up to 15 m. The beach is composed mostly of shell fragments, which make the foreshore much steeper than the fine sand beaches to the north and the south. The backshore is broad and either flat or dips slightly landward where the berm crest is elevated by wave runup. Dunes on North Padre Island are generally stable except for remnants of dune fields that were active during the 1950s drought and blowouts that originate at the Gulf shore and migrate across the barrier. Beach erosion is slow, but persistent on North Padre Island as demonstrated by this sea wall. When the wall was built in the late 1960s, there was a wide beach separating the wall from the Gulf of Mexico. The progradational barriers of the central Texas coast are compound landforms that preserve backbarrier elements of their former transgressive phase, when they were being overwashed and migrating landward. The curvature of channels on the large washover fan on San Jose Island, (>Fig. 1.5.5) depicts the strong influence of hurricanes

Texas

and the counterclockwise flow of overwash generated by the wind. Similar fans also form the backbarrier on Matagorda Island. All these fan-like features are much larger than the washover fans constructed by extant storms, which suggests that storms during the mid to late Holocene may have been more intense. The progradational barrier islands also are characterised by continuous and densely vegetated dunes and broad and relatively stable beaches. Matagorda Peninsula is a low, narrow, migrating barrier that is located on the southwestern flank of the Brazos/ Colorado deltaic headland. It was devastated by Hurricane Carla (1961) that opened numerous washover channels across the barrier. The channels subsequently filled at their Gulf entrance, but they remain as vulnerable sites for future channel incision. Beaches along the Holocene Brazos/Colorado headland, such as at Sargent, Texas, are composed of stiff, organic-rich mud representing former floodplain marsh environments. These rapidly eroding beaches are overlain by thin washover terrace deposits composed of high concentrations of estuarine shells and rock fragments. Considering their overall morphology and welldeveloped beach ridges, Galveston Island and Bolivar Peninsula are progradational barriers. Although Galveston Island is a progradational landform, most of the beaches along the island are eroding. The highest rates of erosion are immediately downdrift (southwest) of the seawall. In the mid 1960s, a wide sandy beach extended seaward of the seawall and a concrete slab at the western end of the seawall formed a ramp for cars to drive directly onto the beach. The beach at this site has retreated more than 200 m and the erosion threatens nearby condominiums. The eastern third of Galveston Island is protected by a seawall that was constructed after the 1900 hurricane killed thousands of people. A field of groynes and riprap were added to protect the wall from persistent beach erosion. In 1995, a beach nourishment project was used to restore the recreational beach in front of the sea wall. The movement of beaches on Bolivar Peninsula is generally systematic with erosion being prevalent on the northeastern end, where it attaches to the Pleistocene headland, relatively stable in the western third, and accreting near the jetties at the tidal inlet entrance to Galveston harbour. The beach accretion is supplied by erosion of updrift beaches. In 1961, Hurricane Carla overtopped the eastern end of Bolivar Peninsula and deposited a wedge of washover sediment as

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much as 1.2 m thick. Subsequent beach erosion has exposed the washover deposits in a backbeach scarp. The headland east of Bolivar Peninsula is composed of stiff Pleistocene mud overlain by a thin wedge of cohesive Holocene marsh mud. A continuous washover terrace composed of sand, shell and rock fragments forms the backbeach. This beach type extends to Sabine Pass, which is the boundary with Louisiana. Beach erosion and frequent flooding have been so severe that the coastal highway is closed. Washover deposit of Hurricane Gilbert (1989) contains boulder size pieces of asphalt ripped up from the roadbed.

References Curray JR (1960) Sediments and history of Holocene transgression, continental shelf, northwest Gulf of Mexico. In: Shepard FP, Phleger, FB, van Andel TJ (eds) Recent sediments, Northwest Gulf of Mexico, American Association of Petroleum Geologists, Tulsa, OK, pp 221–226 Fisk HN (1959) Padre Island and the Laguna Madre flats, coastal south Texas. In: Proceedings 2nd Coastal Geography Conference, Louisiana State University, Baton Rouge, LA, pp 103–151 Hayes MO (1967) Hurricanes as geological agents: case studies of Hurricanes Carla, 1961, and Cindy, 1963. Report of Investigations 61, University of Texas, Bureau of Economic Geology, Austin, TX Kraus NC (2007) Coastal inlets of Texas. USA. In: Proc. Coastal Sediments’07, American Society of Civil Engineers, New Orleans, LA LeBlanc RJ, Hodgson WD (1959) Origin and development of the Texas shoreline. In: Proceedings 2nd Coastal Geography Conference, Louisiana State University, Baton Rouge, LA, pp 197–220 Morton RA (1979) Temporal and spatial variations in shoreline changes, Texas Gulf Coast. J Sed Petrol 49:1101–1111 Morton RA (2002) Factors controlling storm impacts on coastal barriers and beaches – a preliminary basis for real-time forecasting. J Coastal Res 18:486–501 Morton RA, McGowen JH (1980) Modern depositional environments of the Texas Coast. Guidebook 20, University of Texas, Bureau of Economic Geology, Austin, TX Morton RA, Paine JG, Blum MD (2000) Responses of stable bay-margin and barrier-island systems to Holocene sea level highstands, western Gulf of Mexico. J Sed Res 70:478–490 Morton RA, Miller TL, Moore LJ (2004) National assessment of shoreline change: part 1: historical shoreline changes and associated coastal land loss along the U.S. Gulf of Mexico. U.S. Geological Survey Openfile Report 2004–1043, Washington, DC Morton RA, Bernier JC, Barras JA (2006) Evidence of regional subsidence and associated interior wetland loss induced by hydrocarbon production, Gulf coast region, USA. Env Geol 50:261–274 Shepard FP, Phleger FB, van Andel TJ (1960) Recent sediments. Northwest Gulf of Mexico. American Association of Petroleum Geologists, Tulsa, OK

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1.6 Louisiana

Ioannis Georgiou · Mark Kulp · Michael Miner · Duncan FitzGerald

1.  Introduction The Louisiana coastal plain is one of America’s most unique and important areas in terms of regional coastal ecosystems, natural resources, human infrastructure and cultural heritage. Louisiana is located at the terminus of the Mississippi River, North America’s largest drainage basin stretching from the Appalachian Mountains westward to the foothills of the Rocky Mountains, covering a region of more than 3.3 million sq km (>Fig. 1.6.1). The Mississippi River discharges 6.2 × 1011 kg to the Gulf of Mexico annually (Coleman 1988) building a large delta plain (Fig.  1.6.2), creating a landscape dominated by fluvial pathways, wetlands and bays between active distributary networks. Today, the delta plain consists of large expanses of coastal wetlands within a geomorphologic framework of lakes, estuaries and natural levee systems associated with active and abandoned distributaries. Delta plain formation is the result of the delta cycle, which starts with stream avulsion and delta building and ends with abandonment and delta reworking and deterioration by marine processes and subsidence (Colema­net al. 1998). The delta cycle can be divided into two phases,the fluvial-dominated regressive phase and the marine-­ dominated transgressive phase (Roberts 1997). Products of this autocyclic process of delta lobe progradation and subsequent abandonment are the transgressive components of the delta plain that include barrier island/tidal inlet systems, inner-shelf sand shoals, tidal channels and interdistributary bays (Roberts 1997).

Louisiana

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⊡⊡ Fig. 1.6.2 Reedswamp spreading at the mouth of a Mississippi River distributary. (Courtesy Geostudies.)

3.  Active Delta Plain The constructional phase of active delta growth is characterised by rapid infilling of lakes and bays followed by seaward progradation of the local coastline and the formation of a thick wedge of terrigenous clastic sediment on the shelf. Modern examples include the lacustrine deltas of the Atchafalaya Basin, the Atchafalaya/Wax Lake bayhead deltas (Roberts 1997) and the Balize “birdfoot” delta lobe. Progradational events begin with fluvial avulsion; stream capture from an active distributary. Major avulsion events are a product of decreasing gradient and hydraulic efficiency as the main channel lengthens by means of deltaic progradation. While many gradient advantages exist along the lower Mississippi floodplain, erodible substrate (e.g., abandoned channel sand) and favourable floodplain topography (e.g., active and abandoned floodplain channels) are also critical factors influencing avulsions (Aslan et al. 2005). If these criteria are met, a major river flood or hurricane may trigger avulsion, which results in stream capture by a previously less significant distributary and a delta-switching event, causing a gradual abandonment of the formerly active delta lobe in favour of the more efficient route. Subsequent to deltaic abandonment, previously active delta lobes become erosional headlands, and subsidence and marine reworking results in the landward migration of the shoreline (>Fig. 1.6.3). Sediment within the headland is reworked laterally by waves to form barrier islands and ultimately, inner-shelf shoals. The sandy barrier island

systems along coastal Louisiana constitute an important component of the delta-plain ecosystem because of the habitat they provide, their storm-surge buffering capabilities and their role in maintaining marine and estuarine gradients.

4.  Barataria Bay Barataria Bay is an interdistributary bay that has formed between the Lafourche and modern Placquemines/Balize delta lobe. The wetlands and fronting barrier system comprising this region illustrate well the processes of subsidence dominating abandoned delta plain regions. The Barataria barrier system is experiencing some of the highest relative sea level rise rates in the continental USA. During the past half-century rapid relative sea level rise (0.92 cm/year, Grand Isle tide gauge: 1947–2006) and erosional processes within Barataria Bay have led to substantial wetland loss, converting more than 1,100 sq km of wetlands to open water (16.9 sq km/year) since 1935. Conversion of wetlands to intertidal and subtidal environments is a product of several linked processes including subsidence, marsh front erosion and catastrophic scour during large magnitude hurricanes (FitzGerald et al. 2007). Other factors are also important in wetland loss, such as fluid withdrawal from nearby oil and gas fields (Morton et al. 2005), marsh edge erosion caused by local wind-generated wave erosion and marsh excavation during hurricane impacts (Barras 2006).

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⊡⊡ Fig. 1.6.3 Reworking of sediment by waves on the shore of the Mississippi delta. (Courtesy Geostudies.)

Long-term conversion of wetlands to open water (>125 years) has steadily increased tidal exchange between Barataria Bay and the Gulf of Mexico, resulting in larger inlet tidal prism. Two direct consequences of the increasing tidal discharge are the enlarging tidal inlet geometry and growth of ebb-tidal delta shoals (FitzGerald et al. 2004). During the 1880–2006 period of record the ­cumulative inlet cross section had quadrupled in size coincident with the formation of a new inlet (Pass Abel) and drastic increases in size of Caminada and Quatre Bayou Inlets. It has been shown that wetland loss, bay expansion and increased tidal prisms have strengthened tidal discharges through the tidal inlets and increased the competency of the tidal currents. The inlets have deepened, widened and increasingly transported sand-sized sediment further offshore, which together with relative sea level rise explains the loss of sand from the barrier system and a gradual deterioration of the barrier chain (FitzGerald et al. 2004).

5.  Delta Plain Stratigraphy The stratigraphy of the Holocene Mississippi River delta consists of generally fine-grained sediment deposited within a variety of fluvial, deltaic and coastal depositional environments. These sedimentary bodies formed in res­ ponse to deltaic progradation and abandonment, resulting in an assemblage of overlapping, stacked regressive and transgressive units that consist of unconsolidated sands and muddy sediment (>Fig.  1.6.4). During the waning phases of Holocene sea level rise the early and late Holocene delta plains were deposited as shelf-phase deltas. Trans­sgressive salt marsh, interdistributary bay, lagoon and transgressive  sands (>Fig.  1.6.4, green pattern) overlie regressive deltaic muds, peats and distributary deposits (>Fig. 1.6.4, grey pattern). As the shoreline migrates landward during the transgressive phase, there is a landward shift in paralic environments (brackish marsh, inner-distributary bay, saline marsh and lagoonal barrier). The marine flooding surface of a transgressive

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⊡⊡ Fig. 1.6.4 Idealised dip-oriented stratigraphic cross section of the Holocene Mississippi River delta plain. (With permission from ASCE.)

event is marked by a transition from freshwater marsh, swamp or fluvio-deltaic depositional environments to overlying paralic environments through open marine. The marine flooding surfaces ­separate parasequences within the transgressive systems tract. The Teche Ravine­ ment Surface coincides with the maximum flooding surface and separates the transgressive (early and late Holocene delta plains) from the highstand (modern delta plain) systems tracts. During the present highstand of eustatic sea level, delta switching led to the development of progradational parasequences separated by transgressive ravinement surfaces. The St. Bernard delta complex, which began to prograde about 4,000 year bp, is the first of the high-stand delta complexes (Boyd et al. 1989). Except for the modern Balize (Birdfoot) delta, which is deposited at the shelf edge, the Holocene deltas were restricted to relatively shallow inner-shelf waters. They produced thin (10–30-m thick) deltas of widespread coverage (as much as 15,000 sq km). By contrast, the Balize shelf-edge delta is small in extent but consists of a regressive sedimentary interval nearly 120-m thick (Coleman 1988; Kulp et al. 2002).

6.  Chenier plain The Chenier plain (French chêne) of SW Louisiana extends from Sabine Pass, Texas, to Southwest Point Louisiana, with maximum elevations of about 2–6 m (>Fig. 1.6.5). It consists of multiple, shore-parallel, sand-rich ridges that are perched on and physically separated from one another by relatively finer-grained, clay-rich sediment (Penland and Suter 1989). The Chenier plain evolved during the Holocene as a series of progradational mudflats that were intermittently reworked into sandy or shelly ridges to form the modern Chenier-plain physiography. Numerous cycles of deposition and erosion created alternating ridges separated by marshlands. The mudflat component of the plain has ­been built from fine-grained sediment discharged by the Mississippi River and transported westward. Reworking of the mudflat coincided with periods of non-deposition due to an eastward shift of the Mississippi River deposition centres. During these periods coarse-grained sediment and shells are concentrated along the shore forming Chenier ridges. Ridges become isolated when a new ­supply

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⊡⊡ Fig. 1.6.5 The Chenier plain in southwestern Lousiana (top) showing the distribution of cheniers (dark lines) along the coastal plain, and (bottom) the processes of Chenier plain formation. (Cheniers (dark lines) are from Gould and McFarlan, 1959)

of sediment from the Mississippi River distributaries renews mudflat progradation. Repeated seaward growth and retreat along the Chenier plain is thus a consequence of deltaic deposition farther east and the periodic cessation of sediment supply to the Chenier plain as the deltaic depositional centres become abandoned. Currently, the Atchafalaya River is supplying the Chenier plain with fine sediment and mud and parts of the mudflat coast are presently prograding.

7.  Impact of Rising Sea Level and Storms Currently, southern Louisiana is the site of the highest rates of coastal erosion and wetland loss in the United States, including as much as 75 sq km/year of interior land loss (Barras et al. 2003) and up to 23.1 m/year of coastline erosion for select time intervals. Clearly, southern Louisianal is an ecosystem in crisis. Recent estimates of global sea level rise suggest an average rise of between 0.18 and 0.59 m by the end of this century (IPCC 2007). Added to this projected rise is the high rate of local sea level rise due to subsidence, which undoubtedly will

be disastrous for this low lying area of the northern Gulf of Mexico. Between 2000 and 2008, the Mississippi River delta plain has been the landfall of more than seven hurricanes. Four of these storms brought category 3 and higher wind speeds (>178 km/h) and storm surges (>5 m) to the delta plain. These landfalls have been devastating to this already highly ­fragile system producing extensive marine inundation, widespread barrier island erosion, removal of interior marshlands to create open water areas and degradation of the many flood protection systems that protect the heavily populated delta plain. These effects, in concert with decades of anthropogenically created land loss, have created widespread concerns for the future productivity of the delta-plain ecosystems and the ability of the continually diminishing coastal area to provide the resources necessary to the culture and economy that so heavily rely upon the uniqueness of this region. In response to the predicted drastic changes of the delta plain during the next century, numerous land-loss mitigation and management projects have been implemented and many more are proposed in an effort to protect and even revitalise the system. Successful management

Louisiana

of this transgressive coastal zone requires careful attention to how sediment resources can most efficiently be used so  as to replenish depleted shorelines and how riverine diversions can create localised delta plain growth. Approaches undertaken in the emerging concept of Transgressive Management for the Mississippi River delta plain may establish the guiding principles for dealing with coastal zones undergoing similar changes.

References Aslan A, Autin WJ, Blum MD (2005) Causes of River Avulsion: insights from the late Holocene avulsion history of the Mississippi River, U.S.A. J Sed Res 75:650–664 Barras JA (2006) Land area change in Coastal Louisiana after the 2005 hurricanes: a series of three maps. U.S. Geological Survey Open-File Report 2006–1274 Barras J, Beville S, Britsch D et al (2003) Historical and projected coastal Louisiana land changes: 1978–2050: United States Geological Survey, Open File Report, 03-334, 39p (Revised January 2004) Boyd R, Penland S (1981) Washover of deltaic barriers on the Louisiana Coast. Gulf Coast Assoc Geol Soc Trans 31:243–248 Boyd R, Suter J, Penland S (1989) Relation of sequence stratigraphy to modern sedimentary environments. Geology 17:926–929 Coleman JM (1988) Dynamic changes and processes in the Mississippi River Delta. Geol Soc Am Bull 100:999–1015 Coleman JM, Roberts HH, Stone GW (1998) Mississippi River Delta, an overview. J Coastal Res 14:698–716 FitzGerald DM, Kulp MA, Penland S, Flocks J, Kindinger J (2004) Morphologic and stratigraphic evolution of muddy Ebb-Tidal deltas along a subsiding coast: Barataria Bay, Mississippi River Delta. Sedimentology 51:1157–1178

    1.6

FitzGerald DM, Kulp MA, Huges Z, Georgiou I, Miner MD, Penland S, Howes N (2007) Impacts of rising sea level to backbarrier wetlands, tidal inlets, and barrier islands: Barataria Coast, Louisiana: Coastal Sediments ’07. Proceedings of the 6th international symposium on coastal engineering and science of coastal sediment processes, New Orleans, LA, 1179–1192 May 13–17 Georgiou IY, FitzGerald DM, Stone GW (2005) The Impact of physical processes along the Louisiana Coast. J Coastal Res, Special Issue 44:72–89 Gould HR, McFarlan E Jr (1959) Geologic history of the chenier plain, southwestern Louisiana. Trans Gulf Coast Assoc Geol Soc 9: 261–270 IPCC (2007) Climate change 2007: the physical science basis, summary for policymakers. contributions of working group I to the fourth assessment report of intergovermental panel on climate change Cambridge, UK Kolb CR, Van Lopik JR (1958) Geology of the Mississippi River deltaic plain, southeastern Louisiana, US Army Corps of Engineers Waterways Experiment station, Vicksburg, MS, Technical report No. 3-483 Kulp MA, Howell P, Adiau S, Penland S, Kindinger JL, Williams SJ (2002)  Latest quaternary stratigraphic framework of the Mis­ sissippi  River delta region. Gulf Coast Assoc Geol Soc Trans 52:573–82 Morton RA, Bernier JC, Barras JA, Ferina NF (2005) Rapid subsidence and historical wetland loss in the Mississippi delta plain: likely causes and future implications: U.S. Geological Survey Open-File Report 2005–1216, 115p Penland S, Suter JR (1989) The geomorphology of the Mississippi River chenier plain. Marine Geol 90:231–258 Roberts HH (1997) Dynamic changes of the Holocene Mississippi River Delta Plain: the Delta Cycle. J Coastal Res 13:605–627 Törnqvist TE, Bick SJ, van der Borg K, de Jong AFM (2006) How stable is the Mississippi delta? Geology 34:697–700

67

1.7 Mississippi

Ervin Otvos

1.  Introduction

basin has an area of 22,470 sq km that yields an average discharge of 282 m3/s. The Pascagoula-Leaf watershed, 24,600 sq km in area, has 354 m3/s discharge. Small estuaries occur in the mouths of the Biloxi, Jourdan, and Wolf Rivers, Graveline Bayou, and numerous smaller tidal creeks with minor freshwater discharge. Straddling the Alabama state line on the east are rapidly decaying marshlands that formed when the abandoned Escatawpa River delta filled the head of Grand Bay. With its freshwater source lost, the bay ceased to be a separate estuary. Mississippi Sound, the largest estuarine system, is an unusually wide lagoon between a barrier island chain

The mainland coastline of Mississippi is slightly curved and indented by embayments with small bayhead deltas (Otvos 1981). The hinterland has a series of emerged Quaternary terraces. Late Holocene marshlands and beaches fringe the estuarine shores of the mainland, and an extensive marshland on the SW coast surrounds remants of stranded late Holocene relict islands sev­ eral metres high (Otvos 1998). The largest estuaries are in the lower Pearl and Pascagoula valleys, as well as St. Louis Bay and Biloxi Back Bay. The Pearl drainage

sca

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.7.2 Generalised geological map of the Mississippi coastal plain. N MISSISSIPPI

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⊡⊡ Fig. 1.7.3 Generalised cross-section of the Mississippi coastal plain from the Citronelle Upland across faulted Pleistocene formations to the Late Pleistocene and Holocene formations of the coastal fringe.

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and the mainland. Within Mississippi, the barrier island chain includes Cat, West and East Ship, Horn and Petit Bois islands (>Figs. 1.7.1 and > 1.7.2).

2.  Coastal Terraces Rising as high as 120 m south of Jackson in southern Mississippi, the deeply dissected late Pliocene Citronelle upland surface dominates the coastal plain. Near its southern margin along the mainland coast it stands at 15–25 m (>Fig. 1.7.3). Bordering the Citronelle upland Gulfward are two sizable remnants of the late Pleistocene Montgomery terrace occur at 12–15 m elevation in the central and eastern Mississippi coastal plain (Otvos 1998).

3.  Holocene Coastal Landforms An initially rapid, subsequently decelerating, sea level rise determined coastal development on the northern Gulf (Otvos 2004). Indications of higher than present Holocene

    1.7

sea levels have not yet been reliably identified in Missi­ ssippi. Inundation of the mainland was followed by the gradual infilling of river and creek valleys and widespread wetland development. Marshlands formed in areas of low salinity up to 20 km upstream from the present Pearl River delta front. Brackish marshes were established on the ancient banks of the Biloxi River as the drowning valley became the Biloxi Back Bay. Meandering marsh islands outline the former river channel in the bay centre. The former mainland coast is indicated by the narrow 2–4 m high Magnolia Ridge, composed of dunes landward of the southern Hancock County marsh in SW Mississippi (>Fig. 1.7.4). A bifurcated dune-covered 4 km long narrow sand spit prograded about 2,500 year bp from the Gulfport shore of Belle Fontaine Beach in the southeast.

4. The Barrier Island Chain and Mississippi Sound Westward littoral sand transport from Dauphin Island created a line of elongated narrow sandy shoals, on which

⊡⊡ Fig. 1.7.4 South Hancock Marshland. Point Clear Island is a relict barrier island stranded within marshland. (Courtesy NASA Landsat 7.)

Prairie Surface

Magnolia Ridge

Magnolia Ridge

Lakeshore Bayou Caddy

Pt. Clear Island Campbell Island

SOUTH HANCOCK MARSHLAND

Heron Bay SJP

Mississippi Sound

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Mississippi

⊡⊡ Fig. 1.7.5 Lake Borgne and the relict St Bernard delta lobe, showing the South Hancock marshlands and shoals (SHS). (Courtesy NASA Landsat 7.)

Pleistocene upland

St. Louis Bay

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Pearl River Estuary Lake Pontchartrain

buried barrier islands

Stranded Barrier Islands

south Hancock marshlands

Lake Borgne

ancestors of the present barrier islands aggraded, prograding both westward and gulfward since about 4,600–4,400 years ago (Otvos and Giardino 2004). Two relict islands, Point Clear and Campbell, now partially buried, mark the previous westward continuation of the barrier chain in the present SW Mississippi marshland. Prograding from below Magnolia Ridge, they formed soon after 4,400 years bp (>Figs. 1.7.4 and >1.7.5). Before being stranded and partly

Mississippi Sound

Relict Mississippi-St. Bernard Delta

buried, the chain extended across SW Missis­sippi well into SE Louisiana (Otvos and Giardino 2004). Shoaling was related to the growth of a new Mississippi River sub-delta to the south, which stopped island development west of Cat Island about 3900–3700 years ago. Emergence of the island chain separated the large Mississippi Sound from the Gulf. Eroded remnants of a late Holocene Mississippi River sub-delta form the SW shore of the Sound. Within

Mississippi

    1.7

⊡⊡ Fig. 1.7.6 Central and East Mississippi coast. (Courtesy NASA Landsat 7.)

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Camille Cut Georges Cut

Horn Island Petit Bois Island

West & East Ship Islands

Gulf of Mexico ⊡⊡ Fig. 1.7.7 Beach ridges on Cat Island have been truncated and bordered by a dune­covered barrier spit trending SW to NE. (Courtesy United States Geological Survey.)

Mississippi, the Sound is about 112 km long, 10–22 km wide (>Fig. 1.7.6), and generally 2.0–4.0 m deep. Water depths exceed 6 m in and near tidal channels between the islands. While sandy deposits fringe the mainland and island shores, muddy bottom sediment dominate the lagoon floor. The barrier islands between SW Alabama and SE Louisiana originated by westward longshore drift of sand from a large ebb-tidal delta south of Mobile Bay. From about 4,600–4,400 years bp, the narrow islands emerged through vertical accretion on elongated shoals (Otvos and Carter 2008). Sand between the islands was transmitted by way of the elongated shore-parallel sand shoals and small

ebb-tidal deltas located off inlets. The islands expanded through westward and seaward strandplain growth. Cat Island grew 3 km seaward, but westward progradation dominated the other barrier islands. Erosion of the eastern ends of islands was driven both by fair-weather processes and occasional storms. Wide breaches that formed across narrow low sectors during storms led to the revival of sandy shoals. These processes have temporarily or permanently segmented and eliminated extensive barrier island sectors during the past three centuries (Otvos and Carter 2008). Details of changing patterns of erosion and accretion are given in the electronic version of this chapter. They

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Mississippi

⊡⊡ Fig. 1.7.8 The changing coastline of Petit Bois and Dauphin Islands, 1848–1974.

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⊡⊡ Fig. 1.7.9 Belle Fontaine Beach area. (Courtesy NASA.)

Mississippi

    1.7

⊡⊡ Fig. 1.7.10 Erosion of East Belle Fontaine Bluff during a high spring tide in 1987. The dark material on the shore is humate from the Gulfport Formation.

include truncated beach ridge systems and successive ­outgrowths (>Figs. 1.7.7–1.7.10), the growth and destruction of deltas, and a summary of the historical development of the ­barrier islands.

5.  Hurricane Impacts With all hurricane categories considered, the average recurrence interval within 155 km of the Mississippi shoreline averaged 9–10 years. Hurricanes of Category 3 and above recurred once every 32–36 years on average. About 48 hurricanes and tropical storms have had at least a minor impact on the Mississippi coast since 1838. They included Category 3 cyclones Camille (1969) and Katrina (2005). Camille produced a maximum surge of 6.8 m at landfall in the west, while Katrina generated high flood levels in the western islands and on the adjacent mainland coast, reaching unprecedented maxima

of 10–12 and 7–8 m, respectively. They caused rapid erosion and overwash, notably on Cat and Horn islands (Otvos and Carter 2008). Beach erosion along the mainland coast has resulted in periodic beach nourishment.

References Otvos EG (1981) Barrier island formation through nearshore ­aggradation – stratigraphic and field evidence. Mar Geol 43:195–243 Otvos EG (1998) Citronelle Formation, northeastern Gulf coastal plain: Pliocene stratigraphic framework and age issues. Trans Gulf Coast Assoc Geol Soc 48:321–333 Otvos EG (2004) Holocene Gulf levels: recognition issues and an updated sea-level curve. J Coastal Res 20:680–699 Otvos EG, Carter GA (2008) Hurricane degradation-barrier development cycles, northeastern Gulf of Mexico: landform evolution and island chain history. J Coastal Res 24:463–478 Otvos EG, Giardino MM (2004) Interlinked barrier chain and delta lobe development, northern Gulf of Mexico. Sed Geol 169:47–73

75

1.8 Alabama

Ervin Otvos

1.  Introduction Alabama has an embayed mainland coast with broad lagoons, including Mobile Bay and Perdido Bay, behind a sandy barrier island (Dauphin Island) that continues eastward as the Morgan Peninsula, extending to the Florida border (>Fig. 1.8.1). Much of the coastal plain is occupied by the dissected Upland Surface, underlain by late Pliocene deposits. Bluff-bounded, deeply incised wide valleys in the Mobile-Tensaw River system are associated with extensive valley-floor wetlands and a large marshy delta at the head of Mobile Bay. The narrow Holocene coastal plain east of Mobile Bay includes a beach ridge complex, interspersed with narrow lagoons, inlets, paralic lakes and marshes. The outer coast is bordered by a sandy barrier beach, 20 km long and 2–4 km wide, with dunes on the Morgan Peninsula. A large compound barrier spit at the mouth of the bay has narrowed the entrance to Mobile Bay.

In addition to Mobile Bay, there are several smaller estuaries, including Perdido Bay, Fowl River, Fish RiverWeeks Bay and Bayou la Batre. On the inner side of the entrance to Mobile Bay is a large cuspate sandy ebb-tidal delta. To the west, an extension of Mississippi Sound lies between the Dauphin barrier island and the mainland coast.

2.  Coastal Plain The late Pliocene Citronelle depositional upland surface is the oldest and most dominant element of the coastal plain. It is primarily alluvial, but portions were deposited under estuarine conditions (Otvos 2004). It is thought to have been raised by post-Pliocene tectonic uplift. Hundreds of small circular and elliptical depressions, probably of aeolian origin, dot this upland surface.

⊡⊡ Fig. 1.8.1 Major features on the Alabama coast.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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An erosional scarp forms the boundary between the Citronelle upland and the low, gently seaward sloping Prairie alluvial surface. The Prairie plain widens along the western side of Mobile Bay. It includes remnants of Pleistocene barrier islands.

3.  Holocene Evolution Dauphin Island and the adjacent Mississippi barrier chain formed, when the rising sea encroached on and surrounded an isolated elevated remnant of the Gulfport barrier SW of Mobile Bay. Large sand volumes supplied by westward littoral transport from the adjacent large ebb tidal delta built an apron around this barrier remnant.

A line of shoals formed in the nearshore zone and barrier islands were built on them by vertical aggradation (Otvos and Carter 2008). The rising late Quaternary marine transgression gradually drowned the Mobile Basin and smaller valley mouths to the east. Barrier islands then formed (Otvos 2005), enclosing wide lagoons. Morgan Peninsula consists of two obliquely aligned beach ridge, sets overlain by a substantial dune cover (Rodriguez and Meyer 2006). Gradually, constricting the wide entrance to the Bay in the late Holocene (>Fig. 1.8.2), the Morgan barrier spit has slowly extended westward. It has also prograded, forming a series of lobate forelands bearing successively formed beach ridges that have been truncated by subsequent marine erosion (>Fig. 1.8.3).

⊡⊡ Fig. 1.8.2 Mobile Bay and adjacent areas. (Reproduced with permission from Landsat 7.)

Alabama

1.8

⊡⊡ Fig. 1.8.3 Holocene beach ridges in the Morgan-Perdido area.

4.  Mobile Bay The Mobile Bay drainage basin (115,513 sq km) contains the fourth largest river system in the United States. Fluvial discharge from the Mobile is 2,246 m3/s. Bell-shaped Mobile Bay has a surface area of 1,070 sq km and is 50 km long, 12 km wide in the north, 38 km wide in the south. On the northeast, the bay is flanked by 20–30 m high Pliocene bluffs capped by silty sands of the Citronelle Formation that forms the upland surface. The late Pleistocene Prairie terrace alluvium and Holocene deposits flank the bay shore on the SW, W, and SE. Estuarine and fresh water deposits filled a 64 km long stretch of the incised valley, north of the present Mobile River delta. Oyster reefs were abundant. Salt content ranges between 0.9 and 2.8% in the Mobile Bay, and the eastern extension of Mississippi Sound. Microtidal ranges of 0.4–0.7 m prevail, and ebb-tide currents in the deep, narrow inlet channel attain 5.5 km/h. Construction of numerous large reservoirs on tributary streams reduced fluvial sediment transport to the Bay in the twentieth century. 4.8 million tons of suspended and bottom sediment now reach the delta yearly. About 45% of the fluvial sediment discharged, passes through the Delta, but remains in Mobile Bay, while 30% enters the Gulf of Mexico by the deep main tidal pass and Mississippi Sound by way of Grants (Heron) Pass. The 5.4 m deep outlet channel in the southwestern Bay deepens to 17.5 m in the main tidal pass toward the Gulf. Extensive dredging and deposition of sediment since the second half of the nineteenth century has led to radical changes in bottom topography. There have been many changes in and around the entrance to Mobile Bay and on the associated ebb-tide delta. During Hurricane Frederic (1979) small islands were drastically reduced or completely eliminated. Re-emerging from sand shoals, the small islands expand quickly in storm-free years (>Fig. 1.8.4). Details of historical changes on Dauphin Island are given in the electronic version of this chapter. There have been erosive episodes of variable duration on the southern beaches (>Fig.  1.8.5), while two small cuspate forelands prograded also on the sheltered southeastern island shore (>Fig. 1.8.6). The wave energy shadow created by the huge Mobile Pass ebb-tidal delta, affords the shore sector with limited protection. Sets of multiple sand bars parallel the Gulf shoreline in the nearshore zone. Perdido Bay, a drowned valley, is a 7 km long, maximum 6.5 km wide estuary on the Florida state border. Its average depth is 2.4 m. Clayey silt and silty clay cover the bottom areas, and well-sorted clean sands fringe the bay shores.

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⊡⊡ Fig. 1.8.4 Changes in islands and shoal configuration on the Mobile Bay ebb-tidal delta between 1908 (top) and 1936 (bottom).

Alabama

1.8

⊡⊡ Fig. 1.8.5 Localised erosion at Fishing Pier, east Dauphin Island. Erosion due to changes in wave sheltering and sand transport, influenced by shifting positions of Sand Island, just offshore and on the western flank of the Mobile ebb-tidal delta. Remnants of pine forest previously buried by dunes, then exposed by 1993–1994 coastline retreat. The boardwalk in the foreground now leads nowhere.

⊡⊡ Fig. 1.8.6 Two small cuspate forelands on the southeastern coast of Dauphin Island, prograded in the wave-shadow of an adjacent ebb-tidal delta. (Courtesy Scott L. Douglass.)

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References Otvos EG (2004) Lithofacies and depositional environments of the Pliocene Citronelle formation, Gulf of Mexico coastal plain. Southeast Geol 20:1–20 Otvos EG (2005) Coastal barriers, Gulf of Mexico: Holocene evolution and chronology. J Coastal Res SI-42:141–163

Otvos EG, Carter GA (2008) Hurricane degradation-barrier development cycles, northeastern Gulf of Mexico: Landform evolution and island chain history. J Coastal Res 24:463–478 Rodriguez AB, Meyer CT (2006) Sea-level variations during the Holocene deduced from the morphologic and stratigraphic evolution of Morgan Peninsula, Alabama, U.S.A. J Sed Res 76:257–269

1.9  Florida

Charles Finkl

1.  Introduction The nature of the Florida coastline is about 3,663 km long. Its landforms are closely tied to conditions on the continental shelves off the Gulf and Atlantic and coasts. These can be divided into six shelf zones, each associated with energy segments and sediment associations that influence coastal configuration and the character of the shore. Tanner (1985) noted that wave energy was generally low on the Gulf coast (mean breaker height Fig. 1.9.1). The coastline of Florida is well known for its exten­­ sive  sandy beaches, which have a total length of almost 1,300  km. These beaches provide protection from storm surges along developed shores, present valuable habitat and contribute to the economic prosperity of tourism. The beach-dune system is thus a critical component of state infrastructure that requires periodic maintenance through beach renourishment. Florida beaches range from small pocket beaches in the Florida Keys to long unbroken segments that stretch for many kilometres. Navarre Beach is the longest continuous stretch of protected beach (13 km in the Gulf Islands National Seashore) in northwest Florida. Panama City Beach and Miami Beach are among the longest renourished beaches in Florida. According to the morphodynamic beach classification system Florida Atlantic beaches exemplify dissipative and intermediate beach states. Reflective beach states occur immediately downdrift (southern margins) of stabilised inlets, but are of limited extent, rarely spanning more than a kilometre or two of coastline. Gulf coast beaches tend to be dissipative, except along the Panhandle where more intermediate beach states occur. Florida beaches show compositional variation from pure quartz to carbonate with a complete range of intergrades between these. In general, beaches north of the 28th parallel on both coasts tend to be quartzose while

those to the south are of carbonate sand. Shells and shell hash are common to all beaches in Florida, regardless of the sand-size composition. Berm grain sizes tend to range from 0.25 to about 0.35 mm on Atlantic coast beaches (>Fig.  1.9.2) whereas Gulf coast beaches show both coarser and finer grain sizes depending on the amount of contained shells. Fine sand beaches with particle size distributions commonly falling in the range of 0.15–0.20 mm are often referred to as sugar sand beaches. These find sand beaches are considered excellent for sand sculpting. Along some parts of the beach, usually on the back berm near the dunes, the sand is so clean that when dry it squeaks when walked upon. Coarser particle sizes may occur under specialised conditions where there are many shells or where there is a concentration of lag deposits on eroded remains of beaches when the finer grains have been washed away. Berms on these coarser-grained beaches often show grain sizes in the range of ­0.35–4.0 mm. Most Florida beaches were backed by low dunes, but these have been removed along developed shores. Beach colour ranges from light grey, due to mixtures of shell fragments with varying contents of quartz and carbonate grains, to pure white, as in the case of the west coast sugar sand beaches, and black and dark grey sands near Venice on the southwest coast to yellow in Flagler County. White sand is due to quartz grains whereas the black sand is due to the incorporation of fragmented fossils in the beach. Yellow sand is related to the adsorption of iron oxides by coquina shells that become mixed with the quartz sand. Darker colours can also be associated with organic materials mixed with sand on eroded beaches where the barrier has retreated over marsh sediment, as sometimes occurs along the Florida Panhandle. Beach erosion has been a problem: about 225 km of Florida beaches were classified as critical erosion areas in 1985 (FNDR 1985), while 370 km were critically eroded in 1993 (Clark 1993), the difference being related to increased coastal development, changing environmental conditions, and better assessment techniques. Florida has successful beach nourishment programmes on both Gulf and Atlantic coasts. The Atlantic coast has about 50 nourished areas that together have received

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

1.9

Florida

⊡⊡ Fig. 1.9.1 The coast of Florida, showing breaker zone wave energy levels (see text) and marine terraces at 80, 50 and 35 m above sea level. (Courtesy W.F. Tanner.)

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rocky coasts (wave-cut cliff, wave-cut platform) large offshore shoal, measured in kilometers scattered reefs and patches of vermetid gastropods living coral reef spectacular exposures of sabellariid worm reefs in the intertidal zone scattered sabellariid worm reefs well-developed beach ridge plains of late Holocene age – from west to east: St. Vincent Island, Dog Island, Sanibel Island, Marco Island (in good part destroyed by urbanization), and Cape Canaveral. dune fields the +1.5-m mid-Holocene high-sea-level shoreline growing deltas Pre-Holocene delta, perhaps Sangamon in age point from which the wave-cut platform radiates outward in three directions (E, S, W). beach-ridge plains, pre-Holocene in age, probably Sangamon beach ridge plains, perhaps Sangamon but perhaps older Pleistocene old coalesced sets of beach ridges, poorly preserved, perhaps Pliocene in age

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⊡⊡ Fig. 1.9.2 Trend of composite beach grain sizes for the east coast of Florida, showing a general increase in grain size from north (Volusia County) to south (Miami-Dade County). Moving average trends smooth out station variations and the red linear regression line shows the overall longshore trend.

Florida

up  to  65 million m3 of sediment since the mid 1940s. Miami Beach, built from 1978–982, was the largest single construction event in the history of beach nourishment on the U.S. East Coast with about 9.2 million m3 of sediment dredged from several different borrow areas located in inter-reefal sediment troughs. The success of the Miami Beach project may be attributed to the long extent of the nourished area (17 km) that reduced fill spreading rates, relatively low wave energy, and the fact that longshore drifting ends at a very long downdrift jetty where the sand accumulates. Some coastal segments such as in the Florida Keys, are predominantly rocky with small pocket beaches that are only occasionally renourished with sand from upland sources. Still other coastal areas such as the Big Bend (Taylor, Dixie, Levy counties) are mostly marshlands or tropical mangroves as along the southwest shore of mainland Monroe County on the southern tip of the peninsula. Rocky outcrops are rare, but have a strong influence on coastline configuration. This is more obvious on the Atlantic coast. where barrier islands are related to bedrock highs, where the Cape Canaveral cuspate foreland is determined by bedrock shoals offshore, where coral reef tracts follow bedrock lineaments, and where beaches are perched on older marine benches and platforms that were formed during higher Pleistocene sea levels. In northern Palm Beach County and throughout Martin County the Anastasia Formation outcrops as low sea cliffs and shore platforms. Beach profiles are strongly influenced by bedrock outcrops on the seafloor. The shape of surveyed beach profiles along Broward County is determined by the position of the first rock outcrop offshore. Between Miami and Palm Beach Robertson et al. (2008) found that rock outcrops and structural sand flats, as mapped by Finkl et al. (2005), influenced the point beyond which there was no significant sediment transport. Other examples of rock control are associated with reef systems (colonies of worm casts) built by sabellariid worms in surf and subtidal environments along the Atlantic coast.

2.  The Coastline The Gulf coast of Florida between the Alabama state line and Levy County (about 520 km) consists of beaches backed by sandy beach and dune ridges. The western part of this coastal segment is referred to as the Florida Panhandle and the eastern part as the Big Bend. This zone is noted for its white sand barrier island beaches, and stands behind a shelf sector that bears sand ridges, buried channels, and shoreface sand sheets, with numerous

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ebb-tidal deltas and offshore shoals. The Apalachicola delta is bordered by barrier islands with multiple beach ridges indicating stages in growth and truncation. St. Vincent Island, for example, has several sets of beach ridges truncated on its eastern shore. Intermittent progradation has been related to sea level alternations during the past 4,000 years, but has given place to predominant erosion on these shores (Tanner and Stapor 1972). To the east is Apalachee Bay, where wave energy is generally very low because prevailing winds blow from land to sea, the concavity of the coastline provides for divergence of wave orthogonals and hence energy reduction, the sea floor is unusually shallow and wide, so that essentially all deep water wave energy is dissipated by frictional processes in the passage across the shelf. Between the southern margin of the Big Bend in Dixie County and the northwest flank of Florida Bay the coast (about 485 km) encompasses sand ridges and nearshore sand flats. Incident wave energy increases southward from low to moderate (mean breaker height rising to about 50 cm). This coastal segment has barrier islands and lagoons, and offshore there are drowned karstified limestones surmounted by nearshore sand sheets. Ebb-tidal deltas are associated with major estuaries (such as Tampa Bay and Caloosahatchee River). Sand ridges tend to be obliquely oriented to the coast, although shore-parallel and shore transverse ridges occur in restricted locations. Charlotte Harbour is a large embayment protected by a chain of ­barrier islands that extends south to a landward curve at Sanibel Island. To the south the coast is beach-fringed behind Estero Bay, but beyond Naples it becomes irregular, fringed by the Thousand Islands. This is a soft-sediment coastline with mangrove forests that grade inland to estuarine areas and finally Everglades marsh. At Cape Sable the coastline turns eastward behind Florida Bay and the long chain of islands (the Florida Keys) that run out westward to Key West. Florida Bay, encompassing about 2,100 km2, is a shallow (about 1–2 m deep) inner shelf lagoon-like environment, with more than 200 mangrove islands, interconnected basins, and grassy mud banks. The Florida Keys have a rocky Atlantic coastline with low bluffs (up to 2 m high) and pocket beaches. Sand flats and sea grass meadows occur in the channel between the Keys and the Florida Reef Tract offshore. Mangroves occupy sheltered environments, but north of Biscayne Bay the Atlantic coast is fringed by a sandy beach. In this 220 km sector the continental shelf is narrow, being only about 1.5 km wide off central Broward County. The outer edge has a zone of coral reefs that extends northward from

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the Florida Keys. The most prominent geological and geomorphological features along the southeast Florida coast include sand flats, nearshore sand bars and rocky reefs. North of Palm Beach is a 170 km sector of sandy beaches, interrupted by tidal inlets. Some of these have shoals offshore (>Fig. 1.9.3). Southward longshore drifting is indicated by accretion against obstacles (>Fig. 1.9.4) and on the northern side of jetties (>Fig. 1.9.5).

The large Cape Canaveral cuspate foreland has numerous beach ridges indicating stages in its evolution (CraigSmith 2005). Prominent transverse bars and sand ridges round the cape offshore. From Daytona Beach to the Georgian border (about 235 km) sandy beaches are backed by dunes. The continental shelf has large shore-parallel sand ridges, sand waves, sand banks and shoals, with an extensive nearshore ⊡⊡ Fig. 1.9.3 Sebastian Inlet, on the Atlantic coast of Florida. (Courtesy J. Morelock.)

⊡⊡ Fig. 1.9.4 Accretion of southwarddrifting sand alongside a grounded ship at Rivera Beach. (Courtesy J. Morelock.)

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⊡⊡ Fig. 1.9.5 Accretion of southwarddrifting sand alongside a jetty at Port Canaveral. The cuspate foreland Cape Canaveral is in the background. (Courtesy J. Morelock.)

sand sheet. Sand is carried onshore to the beach by constructive swell in fine weather, and withdrawn to the sea floor during storms.

References Clark RR (1993) Beach conditions in Florida: a statewide inventory and  identification of the beach erosion problem areas in Florida.  Tallahassee: Florida Department of Environmental Pro­ tection, Beaches and shores technical and design memorandum, 202, pp 89–91

Craig-Smith SJ (2005) Cuspate forelands. In: Schwartz ML (ed) The Encyclopedia of Coastal Science, Dordrecht, Kluwer, pp 354–355 FDNR (Florida Department of Natural Resources) (1985) Beach restoration. A state initiative. Tallahassee, Division of beaches and shores Finkl CW, Benedet L, Andrews L (2005) Submarine geomorphology of the continental shelf off southeast Florida based on interpretation of airborne laser bathymetry. J Coastal Res 21:1178–1190 Robertson W, Zhang K, Finkl CW, Whitman D (2008) Hydrodynamic and geologic influence of event-dependent depth of closure along the south Florida Atlantic coast. Mar Geol 252:156–165 Tanner WF (1985) Florida. In: Bird ECF, Schwartz ML (eds) The World’s Coastline, Van Nostrand Reinhold, New York, pp 163–167 Tanner WF, Stapor FW (1972) Accelerating crisis in beach erosion. Int Geogr 2:1020–1021

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1.10 Georgia

Miles Hayes

1.  Introduction The coastline of Georgia borders a wide plain on the trailing edge of the North American plate, which is generally tectonically stable to slightly downwarped. The coastline is low-lying and depositional, with Pleistocene and Holo­ cene barrier island, bay, lagoonal and estuarine deposits blanketing an irregular topography eroded during low stands of sea level in the Pleistocene epoch (Hayes 1994). The state borders the western flank of the Georgia Embayment, a large indentation in the coastline of southeastern United States. The outer coastline is about 160 km long, but the intricate tidal shorelines behind the barrier islands are several times that length. Wave heights are generally less than 0.8 m, and mean spring tides are between 1.2 and 2.8 m. The Georgia coast is generally tide-dominated, in contrast with the more wave-dominated coast in Florida to the south, and the strongest currents are generated through tidal inlets. The climate is humid subtropical, with abundant rainfall throughout the year. Hurricanes frequent the area in  late summer and early fall, and extra-tropical cyclonic depressions are common during the winter months. Except during the passage of these cyclonic storms, the winds blow prevailingly offshore, making this a lee coast.

2.  The Coastline The coast of Georgia is dominated by a chain of sandy barrier islands interrupted by sounds and backed by marshy estuarine embayments (>Fig. 1.10.1). Several rivers cross the coastal plain and widen into estuaries that converge behind the barrier islands. The rivers with watersheds restricted to the coastal plain, typically have large estuarine systems at their mouths (e.g., Satillo River), whereas the two rivers that drain the piedmont and mountain regions (Savannah and Altamaha) have created modestsized river deltas. The barrier islands show numerous beach and dune ridges running generally parallel to the Atlantic coastline,

but locally in transverse patterns that have been truncated. The ridges curve back beside inlets past and present. Four of the larger barrier island complexes have younger (Holocene) barrier islands on their seaward sides that are backed by remnants of earlier (Pleistocene) barriers and extensive wetlands. The stratigraphy of the barrier islands shows evidence of phases of progradation and retrogradation: inlets have formed, migrated and been filled (Hubbard et al. 1979). A rising Holocene sea level invaded valleys and lowlands to form the estuarine area, while barrier islands formed from sand either washed in from the continental shelf (having been previously deposited there by rivers) or brought to the shore by the two piedmont rivers. Some of the outer barrier islands have shown intermittent landward migration, notably during hurricane overwash. Transgressive dunes are spilling landward on Cum­ berland Island in the south of the state. Blackbeard Island has a complex pattern of intersecting beach ridges, representing several stages of accretion and erosion, with overall progradation. A marshy corridor, threaded by Blackbeard Creek, links it to Sapelo Island, which has a Pleistocene core, presently without beach ridges. The Holocene outer portion of Ossabow Island, to the north, has grown southward as several recurves. Tybee Island contains marsh with tidal creeks and an outerfringe of beach ridges. Attempts have been made to establish sequences of erosion and accretion in beach ridge systems. While there may be long-term recession of the coastline in response to the rising sea level in some areas, the survival of beach ridges indicates that progradation has occurred within the past 5,000 years elsewhere. With the largest tides in the SE United States, and the flattest coastal plain, the coast of Georgia is essentially inundated twice daily by the tides. The large estuarine embayments are flanked by extensive salt marshes. Howard and Frey (1980) classified the estuaries into two main categories: salt marsh and riverine. Salt marsh estuaries, the more abundant type, are dominated by marine waters. They contain sandy point bars, which are distinguished from fluvial point bars by their more abundant biogenic

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⊡⊡ Fig. 1.10.1 The coast of Georgia (based on a mapby Howard and Frey 1980).

structures and mud content. Riverine estuaries are fed by piedmont or coastal plain rivers. Sediment carried by the piedmont rivers is more immature mineralogically than those of coastal plain rivers, which have passed through more than one cycle of erosion and deposition.

Relationships between organisms and sedimentology are well illustrated by salt marsh zonations, as on Sapelo Island (Edwards and Frey 1977) (>Fig. 1.10.2). Tidal inlets have large ebb-tidal deltas of the type illustrated in (>Fig. 1.10.3).

Georgia  

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⊡⊡ Fig. 1.10.2 Salt marsh zonation on Sapelo Island (after Edwards and Frey 1977).

JUNCUS MARSH

BARREN

DISTICHLIS MARSH

SALICORNIA MARSH

SHORT SPARTINA MARSH

TRANSITIONAL MARSH

MEADOW MARSH

BARREN

MEADOW MARSH

PONDED WATER MARSH

MEADOW MARSH

BACK - LEVEE LOW MARSH

LEVEE MARSH

STREAMSIDE MARSH

CREEK BANK

LAND

HIGH MARSH

LOW MARSH

ca MHWN SANDY MUD

MSL

MUD

MUDDY SAND

SAND

NOT TO SCALE

MHWN - MEAN HIGH - WATER NEAP MEAN SEA LEVEL MSL -

USUALLY ADMIXED WITH SPARTINA

ISLAND

⊡⊡ Fig. 1.10.3 Diagram of ebb-tide delta sediment (After Oertel 1973).

Belt of carbon - rich sands

4 1

5

3 2 6

BARRIE

R

SOUND

References Edwards JM, Frey RW (1977) Substrate characteristics within a Holocene marsh, Sapelo Island, Georgia. Senckenberg. Marit 9:215–259 Hayes MO (1994) Georgia Bight. In: Davis RA Jr (ed) Geol­­o­­gy of the  Holocene barrier island system. Springer-Verlag, Berlin, Chap 7, pp 233–304 Howard JD, Frey RW (1980) Holocene depositional environments of the Georgia coast and continental shelf. Guidebook 20: Excursions

2 3

5

4 1

1 3 1 2

1

Clean sand

2

Clean sand & mud lenses

3

Interbedded sand & mud

4

Mud pebbles in clean sand

5

Interbedded mud & poorly - sorted sand

6

Coarse lag deposits

2 1

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in Southeastern geology; The Archaeology-Geology of the Georgia Coast. Annual Meeting of the Geological Society of America 66–134 Hubbard DK, Oertel G, Nummedal D (1979) The role of wave and tidal currents in the development of tidal-inlet sedimentary structures and sand body geometry: examples from North Carolina, South Carolina, and Georgia. J Sed Petrol 49:1073–1092 Oertel GF (1973) Examination of textures and structures of mud in layered sediments at the entrance of a Georgia tidal inlet. J Sed Petrol 43:33–41

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1.11 South Carolina

Miles Hayes

1.  Introduction The coastline of South Carolina, like that of Georgia, flanks  a wide depositional coastal plain, which includes successive Quaternary formations, mainly sandy barrier islands and swamps, blanketing an irregular topography eroded during low stands of sea level in the Pleistocene epoch (Hayes 1985, 1994). It occupies the NW flank of the Georgia Embayment, a large indentation in the coastline (>Fig. 1.11.1). Coastal barrier islands are intersected and  backed by estuaries and bays, passing landward to freshwater swamps. The outer coastline is about 320 km long. Wave heights are generally between 0.8 m and about 1.2 m at the North Carolina border. Mean spring tides are between 1.5 and 2.8 m, and the morphology of the South Carolina coast is tide dominated, though with important wave effects; it was referred to as a mixed-energy coast by Hayes (1994). The climate is humid subtropical, with abundant rainfall throughout the year. Hurricanes occur, particularly in late summer and early fall, and extra tropical cyclonic depressions cross the coast with some regularity. The prevailing winds are westerly and so offshore.

2.  The Coastline The South Carolina coast is a depositional coastal-plain with barrier islands located between major estuarine systems. Two principle types of barrier islands are present on this coast, those that consistently migrate landward (transgressive) and those that build seaward (prograding). On some parts of the coast the landward-migrating barrier islands may move landward tens of m/year. However, those that build seaward usually grow more slowly, at rates of less than 3 or 4 m per decade. Which of these two types of barrier islands are present at a particular location along the coast depends upon the rate of sea level change relative to the sand supply (Moslow and Heron 1978; Hayes 1994). Diminished, or low, sand supply and/or rapid sea level rise promotes the development of transgressive ­barrier islands,

and the opposite for prograding islands. Which type of barrier island is present is a site-specific issue, with the availability of sand, which comes mostly from both offshore and alongshore sources, being the controlling factor in most cases. There are 36 km of transgressive barrier islands and 120 km of prograding barrier islands on the South Carolina coast. Transgressive barrier islands (>Fig. 1.11.2) are composed of coalescing wash over fans, or a wash over terrace, which is overtopped at high tides, usually several times a year. In the process of migration, the entire wash over terrace complex moves landward, leaving an eroded near shore zone in its wake (>Fig. 1.11.3). As a result of this type of migration, in three dimensions the entire complex consists of a relatively thin (Fig. 1.11.4) are typically composed of multiple, relatively parallel linear ridges of sand topped by fore dunes. As a result of this type of seaward growth, in three dimensions they consist of a wedge of sand 7–10 m thick that has built seaward over offshore mud. Because of the mesotidal setting, the barrier islands are relatively short (Figs.  1.11.5 and > 1.11.6. The ebb-tidal deltas, which contain 77% of the sand in the coastal zone of South Carolina, are giant sediment traps, which markedly influence erosional and depositional patterns on the adjacent barrier islands. They cause updrift progradation and downdrift erosion in some cases (Nummedal et al. 1977). The most notable example of this is where the main channel of a large ebb-tidal switched to the north abruptly in the mid-1850s, trapping the 150,000 m3 of sand/yr that typically moves southward along this coast in a natural groyne effect (Hayes and Kana 1976). At inlets where sand does readily by-pass the inlet, the coastline may build out on the downdrift side, as shown in >Fig. 1.11.6. Detailed studies of the tidal inlets in South Carolina by Hubbard et al. (1979) and FitzGerald (1976) showed that

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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       South Carolina

⊡⊡ Fig. 1.11.1 Location map, South Carolina coast.

the character of the inlets changes from deep and seaward extended (tide dominated) in the south, as at Fripp Inlet (>Fig. 1.11.6), to complex and sand choked (wave dominated) toward the north. In summary, the dominant morphology of the central South Carolina coastline is a mesotidal barrier island complex intersected by numerous tidal inlets and backed by extensive salt marshes and tidal flats. The barrier islands that are prograding have multiple beach ridges, some ­surmounted by dunes, in patterns that indicate complex histories of erosion and accretion, and their persistence

showing overall accretion. Most of the larger, prograding barrier islands have drumstick-shaped configurations, as on Kiawah Island (>Fig. 1.11.7). Charleston stands beside a large estuary with a harbour entrance bordered by spits. To the northeast, Cape Romain is a cuspate foreland built by southward drifting sand, with prograding spits on both sides. The Santee/Pee Dee Delta is the largest river delta on the east coast of the United States, and Winyah Bay is now a large estuary with an entrance deflected southward by a longshore spit. This marks a change to a smoother, arcuate wave-dominated

South Carolina    

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⊡⊡ Fig. 1.11.2 Morphology and subsurface three-dimensional configuration (stratigraphy) of a transgressive barrier island, with a cross-section (A-A’).

⊡⊡ Fig. 1.11.3 A transgressive barrier islands at Edingsville Beach, South Carolina. This coastline has retreated almost a mile since the 1850s. Photograph taken in 1974. (Courtesy M.F. Stephen.)

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⊡⊡ Fig. 1.11.4 Morphology and subsurface three-dimensional configuration (stratigraphy) of a prograding barrier island, with a cross-­sections (A-A’).

⊡⊡ Fig. 1.11.5 Price Inlet, showing ebb tide delta.

South Carolina    

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⊡⊡ Fig. 1.11.6 Fripp Inlet, showing ebb tide delta. The coast to the south (left) has prograded further than that to the north. (Left; downdrift side of inlet).

⊡⊡ Fig. 1.11.7 An oblique infrared view from NE of Kiawah Island, South Carolina, which clearly illustrates the drumstick configuration of this 8.5-mile-long, prograding barrier island. Note the presence of linear ridges of sand vegetated by maritimeforest vegetation (arrow), which indicate the positions of the foredunes just back of the beach at earlier stages in the growth of the island. Photograph taken in 1976: the island is now highly developed with residences. (Courtesy D. K. Hubbard.)

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coastline that extends past Murriells Inlet and Myrtle Beach. The coastline then curves eastward into North Carolina towards Cape Fear. The present coastline shows both eroding and accreting sectors. Continuing changes are evident when recent Google Earth imagery is compared with air photographs published by Shepard and Wanless (1971). Erosion problems have arisen for coastal developers who have built structures too close to unstable and shifting tidal inlets. Captain Sam’s Inlet, located between Kiawah and Sea­ brook Islands, has been migrating southward at the rate of 60–70 m/year (Sexton and Hayes 1983).

References FitzGerald DM (1976) Ebb-tidal delta of Price Inlet, South Carolina. In: Hayes MO, Kana TW (eds) Terrigenous clastic depositional environments; some modern examples, Coastal Research Division, University of South Carolina, Technical Report 11-CRD:143–157

Hayes M  (1985) Atlantic USA – south. In: Bird ECF, Schwartz ML (eds) The World’s Coastline, Van Nostrand Reinhold, New york, pp 207–211 Hayes MO (1994) Georgia bight. In: Davis RA Jr (ed) Geology of the Holocene barrier island system, Springer-Verlag, Berlin, pp 233–304 Hayes MO, Kana TW (1976) Terrigenous clastic depositional environments, some modern examples. Coastal Research Division, University of South Carolina, Technical Report 11-CRD Hubbard DK, Oertel G, Nummedal D (1979) The role of wave and tidal currents in the development of tidal-inlet sedimentary structures and sand body geometry: examples from North Carolina, South Carolina, and Georgia. J Sediment Petrol 49:1073–1092 Moslow TR, Heron SD (1978) Relict inlets: preservation and occurrence in the Holocene stratigraphy of southern Core Banks, North Carolina. J Sediment Petrol 48:1275–1286 Nummedal D, Oertel GF, Hubbard DR, Hine AC (1977) Tidal inlet ­variability – Cape Hatteras to Cape Canaveral. In: Proceedings of American Society of Civil Engineers, Coastal Sediments 77, Charleston, SC, pp 543–562 Sexton WJ, Hayes MO (1983) Natural bar-bypassing of sand at a tidal inlet. In: Proceedings of the 18th International Conference on Coastal Engineering, Cape Town, South Africa Shepard FP, Wanless HR (1971) Our changing coastlines. McGraw Hill, New York

1.12 North Carolina

Stanley R. Riggs

1.  Introduction The North Carolina coastline is a very large and complex system that consists of about 560 km of ocean coast, 23 inlets, over 8,000 km of estuarine shore, with over 8,000 sq km of brackish-water estuaries. North Carolina receives about 1,200 mm of rainfall throughout the year, resulting in a subtropical climate dominated by heavy vegetation. One drainage basin (Roanoke River) flows off the front side of the Appalachian Mountains, four others (Chowan, Tar, Neuse, and Cape Fear rivers) originate within the Piedmont Province, while six others (North, Pasquotank, Perquimins, Alligator, White Oak, and Waccamaw rivers) are small, black-water streams that originate within the Coastal Plain. The coastal system forms the interface between the Atlantic Ocean and a wide coastal plain composed of Cre­ taceous through Holocene sediments. The coastal plain has been formed on the trailing edge of the North American plate, a generally stable tectonic setting with only minor tectonic uplift associated with the mid-­Carolina Platform High (the former Cape Fear Arch) in the southernmost portion of the State. Also, the Albemarle Embayment in the northern portion of North Carolina experienced slight subsidence through the Quaternary. The general geological map of the North Carolina Coastal Plain shown in >Fig. 1.12.1 suggests major differences between the northern and southern coastal regions that reflect their geological heritage – the underlying geological framework. A line drawn from Raleigh through Cape Lookout separates the coastal system into two coastal provinces, each having a unique geological framework that results in distinctive types of continental shelves, barrier islands, inlets, and estuaries. The coastal system in the Southern Province, from Cape Lookout to the South Carolina border, is primarily underlain by rocks of Upper Cretaceous through Pliocene ages. In this region, there is only a thin and highly variable skin of surficial sand and mud of Quaternary age. The older units are generally composed of hard rocks including mudstone, sandstone, and limestone associated with the Carolina Platform, a major geological structure that

underlies the region between Myrtle Beach in South Carolina and Cape Fear in North Carolina. This structural platform rises close to the earth’s surface causing the older and harder rocks to be eroded and truncated by the coastline. This erosional topography results in relatively steep slopes with common exposures of the older rock units that control the coastal geometry. In contrast, the coastal system in the Northern Province, from Cape Lookout north to the Virginia border, is underlain primarily by younger sediment of Quater­ nary age. These generally consist of slightly unconsolidated mud, muddy sand, sand, and peat that thicken northward to fill the subsiding Albemarle Embayment with up to 70 m of sediment. The generally soft sediment were deposited during sea level fluctuations resulting from multiple glaciations and deglaciations of the Quaternary ice ages (Riggs et al. 1992). Consequently, the depositional topography is dominated by young sediment with low slopes and the older rock units are deeply buried beneath these surficial sediments. The different geological frameworks of the Southern and Northern Provinces produce dissimilar sediment supplies and land slopes. The Southern Province has an average subaerial slope landward of the coast of 0.6 m/km. In contrast, the Northern Province has an average subaerial slope of about 0.03 m/km landward of the coast. Thus, the rising sea level floods the disparate slopes producing different kinds of barrier island-inlet systems and associated estuaries. The steeper slopes of the Southern Province produce short, stubby barrier islands with eighteen inlets and narrow back-barrier estuaries. The gentle slopes of the Northern Province produce long barrier islands with only four inlets, and an extensive sequence of drowned-river estuaries that constitute the vast Albemarle-Pamlico estuarine system. These northern barrier islands project seaward, forming Cape Hatteras and the famous Outer Banks: a sand dam that semi-­ isolates the Albemarle-Pamlico estuarine system from the ocean. In addition, the continental shelf morphology within the two coastal provinces is dictated by the underlying geological frameworks. In the Southern Province, Onslow

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⊡⊡ Fig. 1.12.1 Generalized geological map of the North Carolina Coastal Plain showing the two coastal provinces and the four geomorphological compartments of the North Carolina coastal system. The cuspate embayments are defined by the Carolina capes and associated cross-shelf, sand shoals. Due to the different geological framework and spatial geometry, the coastal system within each province and compartment is significantly different from the others. Geological map is modified from the NC Geological Survey State Map (1985).

and Long Bays have broad, shallow shelves dominated by hard bottoms due to the extensive rock-floored character of eroded Cretaceous to Pliocene strata cropping out in a broad sweeping arc around the Carolina Platform. In contrast, the shelf morphology within the two coastal compartments of the Northern Province is narrower, slightly steeper, and dominated by soft Quaternary sediment units that fill the Albemarle Embayment (Riggs et al. 1992). The resulting difference in shelf geometry is a prime factor in determining the type and magnitude of physical energy (waves, currents, and tides) expended on the shoreline, as  well as the geological and biological coastal system response. The North Carolina coastal system consists of four geomorphic compartments, each with its own physical

and chemical dynamics and resulting biological and ­geological components. These compartments, or cuspate embayments, are formed by the capes and associated shoals. Each cape contains a shore-perpendicular, shallow sand shoal that extends seaward across the continental shelf: Diamond Shoals off Cape Hatteras, Lookout Shoals off Cape Lookout and Frying Pan Shoals off Cape Fear. The geographic orientation of each compartment and the variable continental shelf geometry determine wave and current dynamics, astronomical and storm tide characteristics, and the type and response to specific storm systems including the extra-tropical storms (nor’easters) and tropical storms (hurricanes). Consequently the coastal systems within the Northern and Southern Provinces of North Carolina are different.

North Carolina

In the Northern Province, coastal systems consist of very high energy, wave-dominated barrier islands with few inlets  into deeply embayed, shallow, irregularly flooded; wind-tide dominated micro- to nano-tidal estuaries. Whereas, coastal systems of the Southern Province consist of moderate energy, wave-dominated barrier islands with abundant inlets to narrow, regularly flooded, micro- to mesotidal dominated estuaries. Consideration of these morphological differences has significant consequences upon both the deposition and preservation of the Holocene record. In the Southern Province, the shallow depth of the shoreface to inner-shelf ravinement surface has truncated and removed only the tops of the estuarine channel infills leaving a vast network of channels exposed on the continental shelf (Riggs et al. 1995). Within the narrow, back-barrier estuarine systems, the preserved Holocene record generally represents the last 3,000 years and only locally are older (Pleistocene) sediments preserved. Generally poor preservation of the Pleistocene formations is attributed to re-excavation of previously deposited sediment during sea level fluctu­ations. The embayed estuaries of northeastern North Carolina are found within the slightly subsiding Albemarle Embay­ ment, a large Quaternary basin on the mid-Atlantic continental margin that extends southward from the Virginia line to Cape Lookout. This basin contains an enigmatic stratigraphic record displaying an embayment history dominated by multiple episodes of fluvial incisement during glacial events and depositional sequences that infill the incised valleys during subsequent interglacials (Riggs et al. 1992). The infill record reflects general transgressions and regressions with sediment sequences that theoretically grade from fluvial to estuarine, barrier island, and open marine sediment or some combination thereof. Barrier islands show much variation. Many barriers are sediment starved while some have adequate supplies to maintain a healthy beach system (>Fig. 1.12.3). Each barrier segment is dependent upon its inheritance from the gene pool that determines its characteristics and health. The gene pool is a complex interaction between the barriers location relative to potential sediment sources, paleotopography of the underlying geological framework, and physical dynamics of the coastal system (Riggs et al. 1995). This inheritance determines sand supplies, shoreface geometry, barrier morphology, inlet stability and geometry, and shoreline recession rates. Ultimately, sediment availability determines the evolutionary history of each portion of the coastal system. Some barriers have been sediment-rich at various stages in their Holocene evolutionary development if they

1.12

occurred in specific geological settings with significant ­sediment supplies. Three major potential sources of sand exist along the North Carolina coast: (a) Extensive sand and gravel occurs in fluvial channels and deltaic structures deposited by Piedmont-draining trunk rivers on the continental shelf during sea level low stands; (b) Tremen­ dous volumes of sand are potentially available to coastal segments adjacent to cape-shoal structures (Diamond, Lookout, and Frying Pan Shoals); (c) Locally, sand-rich stratigraphic units crop out on the inner continental shelf and shoreface that are being eroded by the ravinement surface. Ultimately, sand availability dictates the character of the barrier island that forms and determines how it subsequently responds to both ocean dynamics on short-term scales and sea level changes on the long-term scale. This in turn dictates the degree of isolation, the physical and chemical characteristics, and ecosystem composition of the sedimentological regime within the back-barrier estuarine basins. Sediment-starved coastal segments develop simple overwash barriers that are eroding landward (>Fig. 1.12.2). Landward migration of a coastline produces a transgressive barrier island that results from either lack of a sediment supply or a rise in sea level. The absence of sand supplies is commonly associated with submarine headlands with shoreface and inner shelf stratigraphy dominated by indurated lithologies or with paleovalleys dominated by estuarine mud infill. Both situations result in negligible production of ‘new’ sand and high shoreline recession rates. Barrier segments with minimal sand supplies are characterised by low and narrow islands dominated by storm-tide overwash processes and frequent opening and closing of shallow and highly migratory inletoutlet systems. Much of the sand is incorporated into storm surge overwash fans and associated flood-tide delta deposits that are often dominated by extensive back-­barrier salt marshes. The resulting islands contain only local and ephemeral dunes with little or no maritime forest or extensive vegetative cover due to the dominance of annual saltwater overwash events. Simple overwash barrier islands are abundant along the North Carolina coast. Many are within the Cape Look­ out National Seashore (including North and South Core Banks) and Cape Hatteras National Seashores (including portions of Pea, Hatteras, and Ocracoke Islands). About 40 km of the latter segments are in a severe state of erosion causing greater difficulty in holding the coastal highway on collapsing barrier island segments. A few overwash islands are in other hands such as Masonboro Island State Park. However, some of these islands are in private ownership and represent areas of severe and ongoing conflicts

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⊡⊡ Fig. 1.12.2 Simple overwash barrier islands form in sand-limited settings. Minimal and ephemeral sand occurs on an eroding shoreface that contributes limited amounts of ‘new’ sediment to the coastal system. These islands are dominated by overwash fans that form during major storm-tide events and produce wide and shallow sand and marsh habitats extending well into the back-barrier estuaries. These shallow flats are quickly colonized by fringing marshes that continue to trap sediment as long as the overwash processes continue unhindered by either natural changes or human development such as building barrier dune ridges, roads, and extensive walls of buildings. If the latter happens, then the back-barrier estuarine shoreline may shift from one dominated by constructive processes to one dominated by loss of fringing marsh habitat and net coastline recession.

between continuing development and natural dynamics. These include Topsail and Figure Eight Islands, Ocean Ise,  Wrightsville and Holden Beaches, and the towns of Rodanthe and Kure Beach. Sediment-rich coastal segments develop complex barrier islands consisting of two or more stacked sand sequences (>Fig. 1.12.3). Multiple accumulations of sediment occurred during previous portions of the Holocene history where and when significant sand supplies were available (>Fig. 1.12.4). If large sediment supplies were available sand accumulated in one or more storage bins that included barrier beach dunes, backbarrier dune fields, or regressive beach ridges (>Fig. 1.12.5). Seaward progradation of a shoreline produces a regressive barrier island that results from either introduction of large sand supplies or a  drop in sea level. If inlet systems develop, sand is stored in ebb- and flood-tide deltas and migrating channel deposits. If an inlet closes down, the ebb-tide delta collapses and supplies sand locally to adjacent barriers. Such

local influxes of ‘new’ sand can result in the development of small-scale backbarrier dune fields. Complex barrier islands are characterised by unique coastal features and are common on the barrier islands. Nags Head Woods, Jockeys Ridge, and Shackleford Banks are characterised by back-barrier dune fields with dunes up to 15 m high (>Fig. 1.12.4), whereas Kitty Hawk Woods, Buxton Woods, and Bogue Banks are characterised by accretionary beach-ridge structures and associated maritime forests (>Fig. 1.12.5). In consort with development of barrier width and height comes the growth of maritime forests and prehistoric and historic human occupation. Complex barriers may also be characterised by large estuarine sand shoal complexes such as Colington Shoals (>Fig. 1.12.4) or the filled portions of back-barrier estuaries such as portions of Currituck, Roanoke, and Core Sounds. If sand supplies decrease through time, barrier segments will grade from progradational to stable to recessional.

North Carolina

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⊡⊡ Fig. 1.12.3 Complex barrier islands form in coastal areas with large and available sand sources. Fluctuations in sea level and ocean dynamics result in accreted beach ridges, welded sand prisms, and development of back-barrier dune fields. These islands are high and wide with extensive sand deposits that prevent storm tides from washing over the top of the island. Thus back-barrier estuarine shorelines have no direct connection with oceanic processes and result in characteristics and processes similar to mainland estuarine shorelines.

Associated beaches respond similarly and grade from those characterised by long-term to short-term accretion and to low annual recession rates. The barrier islands have been much modified by human activites. In northeastern North Carolina the natural processes were forever altered in the late 1930s with construction of barrier dune ridges from the Virginia line south to Ocracoke Inlet. Continued construction and maintenance of the barrier dune ridges for the past seven decades has changed the dynamics of all barrier islands, in particular simple overwash barriers (>Fig. 1.12.6). The barrier dune ridges acted as walls that prevented most overwash processes associated with average storm events. Mini­mizing overwash resulted in little to no sediment delivery to the inner shores of the island long-term maintenance of the bordering marshes. This led directly to increased rates of sound-side ­shoreline erosion (Riggs 2001; Riggs and Ames 2003) and

extensive growth of fresh-water vegetation that characterises the northern barrier islands today. Only large storm events or multiple and sequential smaller storms lead to severe dune ridge erosion that finally breaches the dune, causing development of small overwash fans. The semi-fixed dune ridge, in concert with a natural sand deficiency, and net ocean coastline recession, has caused the ocean beach profile to steepen resulting in ­relatively higher rates of coastline recession. In addition, the Outer Banks that are backed by the vast expanses of highly energetic Pamlico and Albemarle Sounds are experiencing increased rates of back-barrier shoreline recession, causing an overall narrowing of the simple overwash ­barrier segments. Like walls of a wooden fort, the artificial dune ridges significantly altered the dynamic barrier island ­processes. In addition, it provided a false sense of security that encouraged development on the overwash barrier segments and fueled the ever-increasing conflicts

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⊡⊡ Fig. 1.12.4 Schematic map and geological cross section of the Kitty Hawk to Nags Head portion of the NC Outer Banks showing preliminary geological interpretations of a complex barrier island system. Interpretations are based upon high resolution seismic and ground-penetrating radar surveys in consort with limited drilling and radiocarbon age dates. The oldest Holocene deposit is the vast submarine sand shoal called Colington Shoals. It prograded southward over the RoanokeAlbemarle channel complex and encompassed extensive Pleistocene paleotopographic highs. The first sand prism (Kitty Hawk Woods-Colington Island-Nags Head Woods) was deposited by rising sea level on top of Colington Shoals. GPR data suggest that at least one and possibly two younger sand prisms were subsequently welded on to the barrier island.

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North Carolina

⊡⊡ Fig. 1.12.5 Interpretation of beach ridges and beach ridge sets on the Kitty Hawk Woods complex barrier island (from Fisher 1967). The map also indicates the back-barrier dune field that is overriding the beach ridges, and the modern beach prism.

Long poin I B

CURRITUCK SOUND

D

A

C E

SEQUENCE.OF BEACH RIDGE SETS

G

Kitty Hawk Boy

4 3

5

1

6 7

Joo n G uite Bay

2 M

8 9

e field Active dun

Kitty Hawk

Active ld uno fie Defiol

ion

dune Man-mode barrier

plain

Deflation

Kitty Hawk Beach

0

1 mile

N

⊡⊡ Fig. 1.12.6 The 1998 geo-referenced, infrared aerial photograph of a simple overwash- and inlet-dominated barrier island in the area between Avon and Buxton is an example of ongoing “barrier-island narrowing” since the 1852, topographic survey. Since 1852, the ocean shoreline has receded up to 750 m for a net loss of 75% island width in 146 years (average annual erosion rate = −5 m/yr). Also, the coastal highway, built in 1955, has been moved westward four times and now is located adjacent to the back-barrier shoreline and protected from the ocean by a double row of constructed barrier dune-ridges.

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between human and natural dynamics on these geologically ephemeral islands.

Hawk, with a pier along which measurements are made of nearshore profile variations in relation to incident wave and current processes.

2.  The Coastline of North Carolina

References

From the South Carolina border to the Virginia line the coast is entirely low-lying and sandy, dominated by beaches, barriers and backing lagoons and swamps. The chief sites of scientific interest are the cuspate forelands at Cape Fear, Cape Lookout and Cape Hatteras. Cemented shell beds off these capes have modified incident wave patterns and so the shaping of the coastline. There are numerous and variable tidal inlets which have tended to migrate southward: historical maps and charts show how they have changed in position, form and dimensions, some having been sealed off while others have been opened, usually during hurricane surges. There is a major research facility at Duck, to the north of Kitty

Riggs SR (2001) Shoreline erosion in North Carolina Estuaries. North Carolina Sea Grant College Program, Raleigh, Publication UNCSG-O1–11, 69p Riggs SR, Ames DV (2003) Drowning the North Carolina Coast: SeaLevel Rise and Estuarine Dynamics. North Carolina Sea Grant College Program, Raleigh. Publication No. UNC-SG-03-04, 152p Riggs SR, York LL, Wehmiller JF, Snyder SW (1992) Depositional patterns resulting from high-frequency Quaternary sea-level fluctuations in northeastern North Carolina. In: Fletcher CH, Wehmiller JF (eds) Quaternary coasts of the United States: marine and lacustrine systems: SEPM (Society of Sedimentary Geology). Special Publication 48:141–153 Riggs SR, Cleary WJ, Snyder SW (1995) Influence of inherited geologic framework on barrier shoreface morphology and dynamics. Marine Geol 126:213–234

1.13 Atlantic Coast Central (USA)

(Virginia, Maryland, Delaware and New Jersey)

1.  Introduction The coastlines of the Mid-Atlantic Bight in the eastern United States are backed by a coastal plain and fronted by a  broad continental shelf, which is structurally a major ­continental margin geosyncline (the Baltimore Canyon Trough). This shelf is underlain by more than 10,000 m of sediment, the basal deposits dating back to the Triassic. The position of the present coastline in relation to the continental shelf varies from Virginia, where it stands near the outer edge of the continental shelf to northern New Jersey, where it stands close to the hard rock basement. The reason for this variation is probably glaciotectonic. The northern end of the New Jersey coast is very close to the southernmost limit of the North American continental glacial ice sheet in Quaternary times. Many believe that a continental bulge formed, with isostatic upwarping in the area peripheral to the ice sheet at its maximum extent, and that with the waning of the ice sheet and its disappearance from New England, this peripheral bulge subsided rapidly. As the ice sheet melted in Late Pleistocene to early Holocene times, sea level rose toward its present level, penetrating further in the northern area of subsidence, notably into the Hudson River valley, and transgressing less far in Maryland and Virginia. Coastal submergence produced the major estuaries of Chesapeake Bay and Delaware Bay, and minor embayments, fronted by chains of barrier islands. The coast between Virginia and New Jersey is on the track of major storms that pass from south to north: hurricanes, which occur roughly once every 16 years, leaving marker horizons in marsh stratigraphy (Donnelly et al. 2001) and more frequent NE (cyclonic) storms, which occur at random, sometimes as many as five or six a year, and produce wind-driven waves from the northeast. Coastline retreat is sometimes very rapid during ­hurricanes and northeasters, but SE waves arriving across the long fetch of

the Atlantic Ocean generate shoreward ­drifting of sediment into Chesapeake Bay, Delaware Bay and Sandy Hook Bay and seaward drifting when storm waves withdraw sediment from the shore and deposit it on the inner shelf. Tide ranges are small. Mean spring tide range at Cape Henry is 1.0 m, at Atlantic City 1.5 m and Sandy Hook, northern New Jersey, 1.8 m.

2.  The Coastline The coast from the Carolinas to New Jersey is dominated by northward longshore drifting, which has built projecting spits at Cape Henry in Virginia, Cape Henlopen in Delaware, and Sandy Hook in New Jersey. South of each spit was a zone of erosion, which provided sediment, some of which was carried northward by longshore drifting. In Virginia, Maryland and Delaware, the coast intersects uplands of Pleistocene sediment resting on 30–50 m over late Tertiary sediment, while in New Jersey they are composed of Quaternary sediment overlying Cretaceous and Tertiary formations. Supplied with northward-drifting sand, the spits have been shaped by southeasterly waves, refracted around their northern ends. Cape Henry is a lobate foreland with many recurved beach ridges, enclosing a lagoon and marsh area on the southern side of the entrance to Chesapeake Bay. Chesapeake Bay is a large branching ria, formed by submergence of the valleys of the lower Susquehanna River and its tributaries. These were river valleys at the time of the low sea level that accompanied the Wisconsin glaciation, and they became submerged during the late Pleistocene– early Holocene marine transgression. Chesa­peake Bay, the largest inlet on the coastal plain of the Atlantic seaboard, is an estuary 311 km long and up to 40 km wide, opening to the sea between Cape Henry and Cape Charles. On its west coast are the long estua­rine inlets at the mouths of the

Edited version of a chapter by John C. Kraft in The World’s Coastline (1985:213–219). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.13, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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James, Pamunkey, Rappahan­nock, Potomac and Patuxent Rivers, the intervening promontories being locally cliffed, while the east coast is low-lying and marshy, with inlets at the mouths of smaller river valleys and some sandy beaches (> Fig. 1.13.1). The cliffs show erosion by frost action as well as wave attack. Within Chesapeake Bay, there are estuarine meander channels shaped by ebb and flood tidal currents (Ahnert 1960). There is a salinity gradient from sea water (35‰) between Cape Henry and Cape Charles to less than 15‰ at Annapolis and fresh water in the rivers.

Variations occur in response to weather conditions, with freshening by heavy rain and river floods and increasing brackishness in dry seasons. Most of the sediment deposited in Chesapeake Bay have come from the Susquehanna and other inflowing rivers. Deposition of fluvial sediment has increased as the result of soil erosion in the catchment, and has led to the advance of salt marshes, followed by woodland, around river mouths (Froomer 1980). On the other hand, continuing submergence, caused partly by land subsidence, ⊡⊡ Fig. 1.13.1 Sandy beach at Hooper Island, north coast of Chesapeake Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 1.13.2 Waterlogging in subsiding marshland, Hooper Island, north coast of Chesapeake Bay. (Courtesy Geostudies.)

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey)   

has resulted in waterlogging within marshes (> Fig. 1.13.2) and erosion of low-lying islands, some of which have shrunk in size while others have disappeared completely (Hunter 1914). Sharp’s Island in Maryland had an area of 283 ha in the late seventeenth century, reduced by erosion to about 242 ha by 1850. Even in the first decade of the twentieth century, it had farms and a hotel. By 1950, there was just room for half a dozen men to stand on a grassy islet. Now only the Sharp’s Point Light shows where it was (Douglas et al. 2000). The low-lying east coast of Chesapeake Bay borders the Delmarva peninsula, so called because it lies partly in Delaware, partly in Maryland and partly in Virginia. On its Atlantic coast, beaches and barrier islands have formed where sand supplied by longshore and onshore drifting has been delivered to the shore. The beaches are backed by dunes, then a corridor of lagoons and salt marshes, bordering the coastal plain. A scarp lying inland along the southern end of the Delmarva Peninsula may be a relic of an earlier Holocene coastline, formed 3,000–4,000 years ago, but this is difficult to reconcile with a generally accepted sea level history of continuing Holocene marine transgression. Great Machipongo Inlet, Little Machipongo Inlet, Wachapreague Inlet, and Metomkin Inlet are gaps between the barrier islands that have varied in width, depth, and position in response to the contest between tidal currents, which tend to maintain them, and waves, which tend to drift sand onshore and alongshore to divert and sometimes seal them. Several have intertidal and subtidal sandy ⊡⊡ Fig. 1.13.3 The east coast of Cape Henlopen, showing the eroding dunes. A fence has been built in the hope of trapping wind-blown sand to restore the profile. (Courtesy Geostudies.)

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deltas formed by inflowing currents (flood deltas), and more rarely outflowing currents (ebb deltas). At Chin­ coteague, the barrier islands are set forward in front of Chincoteague Bay, a coastal lagoon separated from the sea by elongated Assateague Island, a broad barrier island, which ends in a recurved spit at its southern limit. This is an anomaly on a coast with predominant northward longshore drifting, but probably results from occasional southward drifting during northeasters. At Ocean City, the barrier has been built over, and the beach is backed by a boardwalk. North of Ocean City, the barrier beach abuts the coastal plain, enclosing valley-mouth lagoons, and Rehoboth Beach culminates in the large recurved spit at Cape Henlopen. These barrier islands have been transgressive, moving landward as the result of washovers and blowouts during the continuing slow sea level rise. This is indicated by the stratigraphy of the coastal formations, the barrier sands overlying marsh and lagoon deposits in a vertical sequence that matches the horizontal landward succession. The landward advance of the barriers is accompanied by a landward spread of marshes and lagoons on to the backing coastal plain. Before the arrival of Europeans, Cape Henlopen was similar to Cape Henry in Virginia, a cuspate foreland bordering a marshy lagoon area behind recurved spits that began to form more than 2,000 years ago. After an offshore breakwater was built, the pattern of longshore drifting changed, and a spit began to grow northward in response to SE wave action. Erosion on the east coast (> Fig. 1.13.3)

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is matched by accretion on the north coast. The spit extended by almost a kilometre northward between 1842 and 1977, even though its northward growth has been impeded by the effects of tidal ebb and flow currents. This spit is tending to curve to the northwest, and is likely to develop into a cuspate foreland in the next 300–500 years (Kraft et al. 1980). Delaware Bay is a smaller version of Chesapeake Bay, formed by marine submergence of the lower Delaware valley. Tributary valleys are small, and the bay is simpler in outline, funnel-shaped, 84 km long and widening to 20 km at the entrance between Cape Henlopen and Cape May. The estuarine shores are generally low-lying and marshy, and a channel 10 m deep and 360 m wide has been dredged in the Delaware estuary to give ships access to the port of Philadelphia. From Cape May, a sandy shore curves round to the northeast and becomes another series of barrier islands with intervening gaps, notably Cold Spring Inlet, Hereford Inlet, Townsend Inlet, Carsons Inlet and Great Egg Harbour Inlet, beyond which the sandy barrier widens at Atlantic City. To the north are barrier islands interrupted by Absecon Inlet, Brigantine Inlet, Little Egg Inlet, and Beach Haven Inlet. The barrier then becomes more continuous at Long Beach Island, separated by Barnegat Inlet from Island Beach. As in Virginia, Maryland and Delaware, these barrier islands are backed by an elongated corridor of lagoons with extensive salt marshes, as at Brigantine

(> Fig.  1.13.4), while the inlets have had a history of migration, shallowing, closure, reopening, and deepening in relation to interacting wave and tidal current processes. The sandy coastline shows alternations of erosion and accretion, with major changes in and around tidal inlets (Dolan et al. 1979). Island Beach converges with the mainland coast from Point Pleasant northward to Highlands, where the Sandy Hook spit projects NNW. This is a simple recurved spit, bearing dunes and ridges that indicate stages in its growth. The east coast is exposed to Atlantic waves (> Fig. 1.13.5) and shows northward longshore drifting, while the west coast is generally more sheltered, but exposed to westerly storms. There has been erosion on the eastern shore, countered by armouring (> Fig. 1.13.6) and beach replenishment (Nordstrom and Allen 1980), matched by accretion at the northern end. Human impacts have been extensive on this coast. There are cities such as Rehoboth Beach in Delaware, built on high sectors, and in New Jersey most of the barrier islands are occupied by urban and suburban ­developments, mainly for recreational purposes. Ocean City Maryland, built on a low-lying barrier, is a large city that includes an extensive chain of condominiums. Poor planning and building practices have led to development and building over dunes and on to the backshore in many coastal areas, and as a result northeaster storms and hurricanes have wreaked havoc on this coast. In 1962, a condominium at

⊡⊡ Fig. 1.13.4 Salt marsh at Brigantine. (Courtesy Geostudies.)

Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey)   

   1.13

⊡⊡ Fig. 1.13.5 The beach at Sea Bright, on the east coast of Sandy Hook. (Courtesy Geostudies.)

⊡⊡ Fig. 1.13.6 Western shore of Sandy Hook spit, north of Fort Hancock. (Courtesy Geostudies.)

Ocean City rolled over on its side in a northeaster. Yet, building in this coastal zone is continuing rapidly. Planners are attempting to legislate setback lines to limit the cost of storm destruction, which in the past has come to many billions of dollars, but although the nature of the coastal sedimentary environments, the rising sea level and the

hazards of inevitable storm surges are well understood, there have been few links between planners, engineers, and geologists. This is mainly because the development has been conducted by the private sector, and only recently have people recognised that a serious problem has developed (Cooper et al. 2008).

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References Ahnert F (1960) Estuarine meanders in the Chesapeake Bay area. Geogr Rev 50:390–401 Cooper MJP, Beevers MD, Oppenheimer M (2008) The potential impacts of sea level rise on the coastal region of New Jersey, USA. Climatic Change 90:475–492 Dolan R, Hayden B, Rea C, Heywood J (1979) Shoreline erosion rates along the middle Atlantic coast of the United States. Geology 7:602–606 Donnelly JP, Roll S, Wengren M (2001) Sedimentary evidence of intense hurricane strikes from New Jersey. Geology 29:615–618 Douglas BC, Kearney MS, Leatherman SP (2000) Sea level rise: history

and consequences. Academic Press, New York Froomer N (1980) Morphologic changes in some Chesapeake Bay tidal mar­ ­shes. Zeitschrift für Geomorphologie, Supplementband 34:242–254 Hunter JF (1914) Erosion and sedimentation in Chesapeake Bay. United States Geological Survey Professional Paper 40:7–15 Kraft JC, Allen EA, Belknap DE, John CJ, Maurmeyer EM (1980) Processes and morphologic evolution of an estuarine and coastal barrier system. In: Leatherman SP (ed) Barrier Islands, from the Gulf of St Lawrence to the Gulf of Mexico. Academic Press, New York, pp 149–183 Nordstrom KF, Allen JR (1980) Geomorphologically compatible solutions to beach erosion. Zeitschrift für Geomorphologie, Supplementband 34:142–154

1.14 New York and New England

Henry Bokuniewicz

Introduction The Atlantic Ocean coast of the United States north of 40°30ʹ N is crossed by two important physiographic boundaries. The first is the limit of the coastal plain, a thick sequence of unconsolidated or semi-consolidated Tertiary formations that form a seaward thickening wedge down the eastern seaboard. Long Island (New York State) is the northern limit of this feature, the boundary running south of Staten Island, through The Narrows, up the East River along the east side of Manhattan, down Long Island Sound and out under Block Island, Martha’s Vineyard and Nantucket Island. Some Pliocene sediment, mostly glacially reworked, underlie Cape Cod. Hardly any of this material outcrops, being covered with glacial sediment. Staten Island, to the north of this boundary, is underlain by Triassic and Palaeozoic formations. North of Long Island, New York State, the coastline is principally composed of consolidated bedrock primarily of Cambrian and Ordovician age in the southern reaches (Connecticut) and Devonian in the north (Maine). Notable exceptions are Staten Island, which contains Triassic formations dating back to the formation of grabens at the opening of the Atlantic Ocean, and New Haven Harbour in Connecticut, where the less resistant Triassic sandstones intersect the coast. These rock formations mark the eastern margins of a Triassic graben. The second boundary is the southernmost limit of the  last continental glaciation. The Laurentide Ice Sheet reached these latitudes during the Wisconsin Stage of the Pleistocene Epoch, achieving its maximum extent about 23,200 years bp (Balco et al. 2002). The entire coast has been glaciated, and a thick blanket of glacial drift covers Long Island and Cape Cod. Recession of the ice left a terminal moraine from about 18,000 years bp that crosses Staten Island. It is breached at The Narrows and continues along to divide on either side of Peconic Bay on Long Island. At Lake Success a younger recessional moraine continues along the north shore of Long Island to the ‘north fork’ (Orient Point), and from there to the ocean coast of western Rhode Island and the north shore of Cape

Cod. The older recessional moraine crosses to the south fork (Montauk Point) and its eastern fragments form parts of Block Island, of Nantucket and, of Martha’s Vineyard. Deposits on Martha’s Vineyard have been dated 25,000 and 22,000 years bp. A younger moraine along the shore of Buzzards Bay, Massachusetts was laid down 18,800 years ago (Balco et al. 2002). The south coast of Long Island is fashioned in outwash sands, while the more northern parts exhibit a variety of eskers, drumlins, moraine segments and ground moraines. During the last century sea level has been rising at an average rate of about 3 mm a year along the southern coasts of this region and 2 mm a year in the north, as documented from tide gauge measurements (Emery and Aubrey 1991); radiocarbon evidence suggests that the rate of sea level rise has increased over the past few hundred years (Donnelly 2006). The tides are semi-diurnal. The coastline is microtidal in the south, where ocean tides are less then 1 m along the southern Long Island coast. Tidal ranges are 30 cm or less in restricted embayments, although a resonant standing tide, with a range in excess of 2 m, forms in western Long Island Sound. Tide ranges increase northward to 3 m at Boston and 6.5 m in northeastern Maine. Except for restricted embayments, notably Long Island Sound, Narragansett Bay and Cape Cod Bay, the coastline is exposed to open ocean waves propagated over a wide continental shelf. Significant wave heights can exceed 1 m. The region is subject to hurricanes, tropical storms and nor’easters, which are a combination of high, atmospheric pressure over land paired with low pressure offshore. Nor’easters produce strong winds from the northeast, which drive up water levels along the coast by Ekman transport. So-called storm tides (storm surges) that reach 1 m recur annually, while events with recurrence intervals of a century approach an elevation of nearly 2 m. The south coast of Long Island is divided into three stretches (Taney 1961). The western (ocean) coastline is a series of barrier islands backed by large, shallow bays (lagoons). Tidal wetlands and marsh islands are prevalent, especially in the western bays. All the tidal inlets have been

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.14, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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stabilised. Near the eastern end of the Island the beach is cut directly into a glacial outwash plain in front of the terminal moraine. This section contains several small coastal ponds, which appear to be drowned proglacial stream valleys that have been isolated from the sea by beaches built and maintained by the prevalent, westward longshore drifting. Beaches are typically 30–60 m wide and ­composed of fine to medium quartz sand (>Fig. 1.14.1). The extreme eastern point of the Long Island ocean coastline, Montauk Point, is a high coastal bluff of unconsolidated rock capped by the terminal mo­­raine (Ronkonkoma moraine). The north coast of Long Island on Long Island Sound can be divided into two stretches. The western part is highly embayed as a result of the drowning of an irregular, ground moraine draped over antecedent topography, while the eastern section is a gently curving, uninterrupted line of high, unconsolidated coastal bluffs. The coastal plain is absent from the Connecticut coast, which has been formed by the partial submergence of the Coastal Slope peneplain, cut into Palaeozoic and ­Pre-Cambrian bedrock. The Coastal Slope has relatively low relief, and intersects sea level at a slope of about 1:100.

It is indented, harbouring numerous, small, isolated salt marshes. The Connecticut coast is a mosaic of consolidated bedrock (>Fig.  1.14.2; e.g. between Groton and Ston­ington parts of the Thimble Islands), pocket beaches, tombolos (e.g. Long Beach in Stratford, Con­­necticut) or spits (e.g. Griswold Point), all composed of reworked, glacial sands (>Fig. 1.14.3). Sandy beaches are ­isolated and usually less than 3 km long. Some are cut into sandy outwash heads like the Lordship Outwash at the mouth of the Housatonic River. Other stretches of the Connecticut coastline is composed of moraine fragments (e.g. the Norwalk Islands), and some, coastal bluffs. The western coast of Rhode Island is a series of baymouth barriers between sectors of coastline cut into mo­­ rainic sediment (Fisher 1985). Westward longshore drifting has formed beaches across the mouths of drowned, proglacial outwash stream valleys. The eastern, Rhode Island shore is deeply embayed, as at Narragansett Bay, which lies in a structural basin formed in less resistant Carboniferous slates and schists. Bedrock islands are mantled with glacial drift and sandy sediment is sufficient to form only minor pocket beaches. The recessional moraine is absent across

⊡⊡ Fig. 1.14.1 Southampton, New York. The steep incline of the foreshore is due to the relatively coarse nature of the sand. The foreshore becomes less inclined to the west as the sand becomes finer. (Courtesy Bret Bennington.)

New York and New England

⊡⊡ Fig. 1.14.2 Bedrock coastline at Bluff Point, Connecticut. (Courtesy Nancy McHone and Ralph Lewis.)

⊡⊡ Fig. 1.14.3 Bluffs and beach at Bluff Point, Connecticut. (Courtesy Nancy McHone and Ralph Lewis.)

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the mouth of Narragansett Bay, but reappears as the Elizabeth Islands, which are islands of low relief composed of glacial drift and sheltering Buzzards Bay. Boulders from the moraine form a natural protection for sectors of the coastline. The low-energy coast of Buzzards Bay is cut into till and glaciofluvial deposits. The moraine continues eastward on the north side of southern Cape Cod. The large outer islands - Block Island, Martha’s Vineyard and Nan­ tucket Island are situated between 10 and 40 km from the mainland. Each has a section of coast cut into fragments of the terminal moraine that forms the south fork of Long Island. On Block Island the erosion of the moraine and its associated ground moraine results in high coastal cliffs. Sandy beaches have been formed on the east and west shores by longshore drifting (Fisher 1985). The moraine stretches along the northern parts of Martha’s Vineyard and Nantucket Island, while their southern shores are cut into glacial outwash plains. Cliff recession of up to 100 m year was measured on Wasque Point by Kaye (1973). Spits and bay-mouth barrier beaches were formed from sand eroded from the coastal cliffs. Notable on the north shore of Nantucket Island are a series of six cusps formed on the inner side of the spit enclosing Nantucket Harbour. The formation of these cusps has been ­attributed to reversals in wave direction (Rosen 1975), although other mechanisms such as tidal eddies or seiching have been proposed. The ocean coast of Cape Cod can be divided into four sections. The southern coast is cut into a sandy, pitted outwash plain south of the moraine. Submergence of the

proglacial topography has produced a series of coastal ponds and baymouth barriers. Cape Cod then extends northward, bordering Cape Cod Bay. This stretch of coast marks the boundary between two lobes of the Laurentide ice sheet, and is divided into the other three sections. The middle stretch is one of unconsolidated coastal cliffs, in excess of 50 m in elevation, cut into the glacial sediment (>Fig. 1.14.4). Longshore drifting of sand in front of these bluffs has built large spits both to the north (Provincetown) and south (Nauset and Monomoy Islands). The largest coastal sand dunes in the region are parabolic dunes found at Provincetown along the northern spit, which approach 60 m in elevation. The southern spit appears to undergo a cycle of breaching and inlet migration southward with a period estimated to be about 150 years. A series of post-glacial relict shoreline features have been identified submerged in Cape Cod Bay. The present coastline of Cape Cod Bay has relatively small coastal cliffs along its eastern margin (>Fig. 1.14.5) shaped by storms from the west (Uchupi et al. 2005). The embayed coastline has been reworked into spits and barriers. Boston Harbour sits in a Pre-Cambrian basin, and partly submerged drumlins appear as a scattering of islands around the Harbour, reworked at their margins into spits and tombolos. North of Boston rocky headlands dominate the coastline to Cape Ann, and north of this headland Plum Island (8 km long) is one of the northernmost large barrier islands. The New Hampshire coastline is of low relief. ­Cres­cent  beaches are found between rocky headlands (Fisher 1985; >Fig. 1.14.6). Post-glacial submergence has

⊡⊡ Fig. 1.14.4 Southern limit of unconsolidated coastal bluffs on the Cape Cod coastline. Sand spits extend southward. (Courtesy Jim O’Connell.)

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⊡⊡ Fig. 1.14.5 Beach at East Sandwich on Cape Cod Bay, Massachusetts. (Courtesy Jim O’Connell.)

⊡⊡ Fig. 1.14.6 Shore of weathered, metamorphic rock at Pemaguid Point, Maine. (Courtesy Maine Geological Survey.)

produced estuaries with fairly restricted access to the sea on rocky, glaciated terrain. Great Bay, at the New Hampshire-Maine border, for example, is a large, shallow estuary set in crystalline metamorphic and igneous bedrock covered irregularly with till and glaciofluvial sediment. It is connected to the sea through a narrow channel (the Piscatagua River) running about 15 km to the coast.

About half of the tidal coastline of Maine is composed of coastal cliffs cut into unconsolidated glacial drift. On the remainder, gently sloping granite descends to irregular coastlines, as on Baker Island, often with a scarp where the blanket of glacial drift has been cut back to expose the  crystalline basement rock. Glacial striations are seen on the granite surface, in places extending

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down through the zone of breaking waves. Cliffs are rare on the granite. Where there are steep coastal slopes, they are often inherited from older landscapes rather than of marine origin. The imprint of glaciation is obvious on Mount Desert Island, which is a high ridge crossed by a N-S glacial trough.

The Maine coastline can be sub-divided into four sections (Kelly 1987) dominated by the bedrock ­structure (>Fig. 1.14.7) and geomorphology. The southern most stretch is one of arcurate bays, where most of the sandy beaches are found. This section is composed of northeaststriking metasedimentary rocks blanketed by glacial sands ⊡⊡ Fig. 1.14.7 Jenness Beach, Rye, New Hampshire. (Courtesy Robert Roseen, University of New Hampshire.)

⊡⊡ Fig. 1.14.8 Gravel beach between rocky headlands at Machiasport, Maine. (Courtesy Marine Geological Survey.)

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and mud. Some of Maine’s largest salt marshes have formed behind these beaches (Tanner et al. 2006). Further north the coastline is indented by long, narrow, fjord-like bays that trend perpendicular to the coast, a signature of tightly folded metasedimentary bedrock. The inlets are rias rather than true fiords. Mudflats and salt marshes occupy sheltered embayments. Stretches of both em-bayed coast and straight, rocky cliffs can be found in the ­northernmost section (Tanner et al. 2006). The bays here are broader than those found further south. Beaches tend to be sandy in the southern section of the Maine coast, but shingle pocket beaches can be found to the north (>Fig. 1.14.8).

References Balco G, Stone JOH, Porter SC, Caffee MW (2002) Cosmogenic-nuclide ages for New England coastal moraines, Martha’s Vineyard and Cape Cod. Massachusetts USA. Quat Sci Rev 21:2127–2135 Donnelly JP (2006) A revised Late Holocene sea-level record for northern Massachusetts USA. J Coast Res 22:1051–1061

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Emery KO, Aubrey DG (1991) Sea-level, land-level and tide gauges. Springer-Verlag, New York, 237p Fisher JJ (1985) Atlantic USA – North. In: Bird ECF, Schwartz ML (eds)  The World’s Coastline, Van Nostrand Reinhold, New York, pp 223–234 Kaye C (1973) Map showing changes on shoreline of Martha’s Vineyard during the past 200 years. US Geological Survey Field Studies Map, MF-534 Kelly JT (1987) An inventory of environments and classification of Maine’s estuarine coastline. In: Rosen P, FitzGerald D (eds) A treatise on glaciated coastlines, Academic Press, San Diego, CA, pp 151–176 Rosen PS (1975) Origin and processes of cuspate spit shorelines. In: Cronin LE (ed) Estuarine Research, Vol 2, Academic Press, New York, pp 77–92 Taney NE (1961) Geomorphology of the south shore of Long Island NY. Department of the Army, Beach Erosion Board, Corps of Engineers, Technical Memorandum 128:51 Tanner BR, Perfect E, Kelley JT (2006) Fractal Analysis of Maine’s glaciated shoreline tests established coastal classification scheme. J Coast Res 22:1300–1304 Uchupi E, Giese G, Driscoll N, Aubrey DG (2005) Postglacial geomorphic evolution of a segment of Cape Cod Bay and adjacent Cape Cod, Massachusetts, USA. J Coast Res 21:1085–1106

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1.15 Great Lakes (USA)

Mary-Louise Byrne

1.  Introduction The Great Lakes consist of five interconnected freshwater lakes: Superior, Michigan, Huron, Erie and Ontario (Carter and Haras 1985). Their coastlines, including islands and connecting channels, total about 17,000 km. The lakes occupy tectonic basins modified by glacial erosion and are underlain, and in places bordered by solid rock. Pre-Cambrian igneous and metamorphic rocks of the Canadian Shield underlie Lake Superior and the northern shore of Georgian Bay, whereas Palaeozoic dolomites, limestones, shales, and sandstones underlie the other lakes (Hough 1958). Pleistocene drift deposits, mostly till and lake-related glacial deposits, mantle the rocks over much of the basin. Lake shore rock outcrops and drift deposits are commonly fronted by narrow (less than 15 m wide) sand or shingle beaches. Although, the overall physical characteristics of the basin were probably formed before the Pleistocene glaciation – there are earlier shales and limestones beneath the basins of Lakes Michigan, Huron, Erie and Ontario – the Pleistocene epoch left the most visible legacy in the form of the unconsolidated drift deposits on the floors of the lakes and around their shores. The lakes had various forms during the melting and retreat of the Wisconsin ice sheet after 18,000 years bp, at first draining southward to the Mississippi, but as the ice completely melted they took up their present positions. Postglacial isostatic rebound also had a profound effect on the basin, notably with respect to former lake elevations. The coastlines of Lakes Huron and Ontario have been rising relative to those of Lake Erie and  Lake Michigan by about 0.3 m/century. Larson and Schaetzl (2001) summarized the geological evolution of the Great Lakes, stating that issues with the modern Great Lakes are concerned with the shoreline configurations and changes in water level. Because the Great Lakes are located in the interior of the continent, between the source regions of the polar and tropical air masses, the basin has complex, rapidly

changing weather patterns. However, the lakes exert a modifying effect on temperature and humidity by acting as heat sources or sinks. Mean annual precipitation ranges from 746 mm (559 mm rain and 187 mm as snow) in Thunder Bay in the northeast of Lake Superior to 893 mm (804 mm rain and 89 mm as snow) in Kingsville on the south of Lake Erie and 964 mm (791 mm rain and 183 mm as snow) in Kingston on the southeast of Lake Ontario. The prevailing winds come from the west, and are the main cause of surface waves and currents. They can also produce short-term fluctuations in lake level during and after intense storms; fluctuations as great as 5 m on Lake Erie. Because the lakes are small, astronomic tides are limited, usually less than 5 cm in range; but seasonal fluctuations, related largely to variations in precipitation and evaporation, cause a seasonal variation of about 35 cm. Prolonged periods (several years) of abnormal precipitation or evaporation result in long-term oscillations of up to 2 m. Erosion and accretion on the lake shores is related primarily to these fluctuations (U.S. Army Engineer Division 1971; Larsen 1973). The lakes were a well-regulated natural system because of their small connecting channels, but two of them, Superior and Ontario, are now artificially regulated. There has been extensive urban and industrial development along the shores of the lakes, particularly near Cleveland and Chicago. Lake Superior is the largest and the most elevated (183 m above sea level) of the Great Lakes. Its regulated outflow drops 7 m in its journey of about 60 km down the St. Mary’s River before spilling into Lake Huron. Lake Huron is connected to Lake Michigan by the wide Straits of Mackinac and the two lakes are a hydrological unit, both being 176 m above sea level. The St. Clair River flows southward to Lake St. Clair and the Detroit River continues into Lake Erie, the fall in water level being only 2.5 m over a distance of about 120 km. At the eastern end of Lake Erie the Niagara River flows northward to Lake Ontario. It is  just over 50 km long and has an average discharge

A revised version of a chapter by C.H. Carter and W.S. Haras (1985).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.15, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 1.15.1 The cliff at Niagara Falls. (Courtesy Geostudies.)

of 57,000 m3/s over the Niagara Escarpment (>Fig. 1.15.1), falling about 50 m to create the spectacular Niagara Falls. The remaining 50 m difference between the levels of the two lakes is spread over 10 km of a much ­narrower river channel below the falls, where a ­thundering current rushes through a deep gorge before it dies out in Lake Ontario, 74  m above sea level. St. Lawrence River transports the outflow of Lake Ontario some 800 km downstream to the Atlantic Ocean.

2.  Lake Superior The American coastline of Lake Superior is about 2,990 km long, including islands. The southern coast consists mostly of Pre-Cambrian rock, but the relief is lower (less than 60 m) and more irregular. In addition to the rock are clay banks, sand and gravel banks, and dunes and marshlands. Sand and cobble beaches alternate with rocky stretches without beaches. Sites of special interest along the southern shore include Pictured Rocks National Lakeshore and the Grand Sable sand dunes, Quaternary sands that rise to about 60 m above the lake. The Apostle Islands National Lakeshore lies near the western end of the lake. It is now thought that Lake Superior separated from Lake Michigan about 1,200 years ago (Johnston et al. 2007), on the basis of the evidence of regional tilting indicated by water level gauges (Mainville and Craymer 2005).

3.  Lake Michigan Lake Michigan, the only lake to lie entirely in the United States, has a coastline length of about 2,635 km. The NW coast, which includes Green Bay, consists of wetlands, limestone bluffs up to 60 m high, and coastal slopes made up of Pleistocene drift, fronted by narrow sand and gravel beaches derived from the eroding glacial drift. Further south, along the western coast, the overall relief of the till cliffs increases to 40 m high, again with narrow sand and gravel beaches. Towards the southern end of the lake, the coast is largely artificial from Chicago, to Gary, Indiana where beaches are angled between groynes but the beaches widen along the southeastern shore. Sites of scientific interest include the Indiana Dunes National Lakeshore and the Sleeping Bear Dunes National Lakeshore. North­ward along the east coast, sand beaches are backed by dunes up to 135 m high, separated by bluffs and slopes in glacial drift. A complex history of lake level changes is indicated by the various types of dunes on the Michigan coastal plain (Arbogast et al. 2002; Timmons et al. 2007). There have been episodes of intermittent beach progradation, marked by the evolution of parallel dune and beach ridges, during phases of falling lake level (Olson 1958), but these are cut back when the lake level rises (Larsen 1973): an illustration of the Bruun Rule (Schwartz 1967). Baedke and Thompson (2000) analysed the relationship between isostasy and lake level and found that there have been ­different rates of vertical movement

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caused by isostasy and/or tectonism around the lake even though there are similar lake-level variations. The NE coast, including the branching inlets of Grand Traverse Bay, consists in part of marshes, sand and clay banks in addition to bluffs of ­glacial drift.

into Lake St. Clair. This small lake, which acts as part of the connecting channels between Huron and Erie, has a coastline of about 272 km; Lake St. Clair is bordered by wetlands. Anthropogenic structures dominate the western side of the lake, beside the city of Detroit.

4.  Lake Huron

6.  Lake Erie

Lake Huron has a coastline length on the American side of about 4,111 km, including 932 km on islands, which are more numerous than those in the other Great Lakes. Saginaw Bay is inset between Au Sable Point and Point aux Barques on the west coast. Between Mackinac and Saginaw Bay, the west coast of the lake is generally low-lying, with marshes and low slopes of glacial drift or bedrock. South from Point aux Barques there are cliffs and bluffs in glacial drift up to 12  m high. Except for the marshes, most of the shore is fronted by narrow sand beaches that increase in width toward the southern end of the lake.

Lake Erie, the shallowest and southernmost of the Great Lakes, has an American coastline of about 808 km, including 69 km on islands. Along the south coast the relief is low, and the shore is made up of sandy beaches and barriers alternating with cliffs and bluffs in glacial drift, bedrock, and some marshy sectors. The coast has been modified by port and industrial development at Toledo and Cleveland. There are some low limestone islands north of Sandusky, near the western end of the lake. The long, fairly straight coast east of Cleveland has cliffs and bluffs of drift and shale up to 30 m high, and eastward longshore drifting has built a large recurved sand spit that runs out at Presque Ile to shelter the port of Erie. Erosion along its western shore has been countered by beach nourishment and the building of a chain of nearshore breakwaters (>Fig.  1.15.2). To the east, the coast is made up largely of shale bluffs commonly capped by glacial drift.

5.  Lake St. Clair At its southern end, Lake Huron flows into St. Clair River in a low-lying corridor that opens through a marshy delta

⊡⊡ Fig. 1.15.2 The west coast of Presque Ile spit, Lake Erie, showing nearshore breakwaters. The sand lobe in the foreground is in the lee of a breakwater. (Courtesy Geostudies.)

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7.  Lake Ontario The American coastline of Lake Ontario is about 608 km long, with 88 km on islands. It has an average depth of about 86 m, with a maximum depth of 244 m. Along the south coast of the lake, from the western end are cliffs and bluffs up to 18 m high composed of glacial drift and outwash deposits. Further east, the bluffs decrease in size, and sandy beaches and barriers appear with some marshes. At the eastern end of the south coast, there are glacial drift cliffs up to 21 m high, sand barriers, dunes and marshes. The beaches are generally narrow and composed of sand  and gravel. A curved sandy beach forms the east coast, and to the north are numerous islands and bays in the approach to the Thousand Islands Channel that leads into St. Lawrence River.

References Arbogast A, Hansen E, Van Oort M (2002) Reconstructing the geomorphic evolution of large coastal dunes along the southeastern shore of Lake Michigan. Geomorphology 46:241–255

Baedke S, Thompson T (2000) A 4,700 year lake level and isostasy for Lake Michigan. J Great Lakes Res 26:416–426 Carter CH, Haras WS (1985) Great Lakes. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 253–260 Hough JL (1958) Geology of the Great Lakes. University of Illinois Press, Urbana, IL Johnston JW,   Thompson TA, Wilcox DA, Baedke SJ (2007) Geomor­ phic and sedimentologic evidence for the separation of Lake Superior from Lake Michigan and Huron. J Paleolimnol 37:349–364. DOI 10.1007/s10933-006-9052-3 Larsen C (1973) Variation in bluff recession in relation to lake level fluctuations. Illinois Institute of Environmental Quality, Chicago, IL Larson G,   Schaetzl R (2001) Origin and evolution of the Great Lakes. J Great Lakes Res 27(4):518–546 Mainville A, Craymer M (2005) Present-day tilting of the Great Lakes region based on water level gauges. Geol Soc Am Bull 117: 1070–1080 Olson JS (1958) Lake Michigan dune development. J Geol 66:254–263, 345–351 and 473–483 Schwartz ML (1967) The Bruun theory of sea level rise as a cause of shore erosion. J Geol 75:76–92 Timmons EA, Fisher TG, Hansen EC, Eisaman E, Daly, T, Kashgarian M (2007) Elucidating aeolian dune history from lacustrine sand records in the Lake Michigan Coastal Zone. The Holocene 17:789–801 U.S. Army Engineer Division, North Central (1971) National Shoreline Study, Great Lakes Region Inventory Report. United States 93rd Congress, 1st Session, House Document 93–121, 5:1–221

1.16 Hawaii

Charles H. Fletcher · Eden J. Feirstein

1.  Introduction The state of Hawaii consists of eight main islands: Hawaii, Maui, Kahoolawe, Lanai, Molokai, Oahu, Kauai, Niihau, and 124 small volcanic and carbonate islets offshore of the main islands, and to the northwest (> Fig. 1.16.1). The climate is sub tropical oceanic. Honolulu has a mean monthly temperature of 21.7°C in January, rising to 25.6°C in July, and an average annual rainfall of 802 mm. The Hawaiian archipelago lies in the zone of northeast­­erly  trade winds, which bring high rainfall to the windward mountain slopes while the leeward coasts are in rain shadow. Hawaii’s preeminent example of this orographic rain is Kauai’s Mount Waialeale, which receives an average 1168 cm of rain a year. In contrast, the dry air descending down Kauai’s leeward

side creates local semi-arid conditions at Polihale Beach on the west side of the island which receives on average a mere 20 cm of rain a year. In addition to northeasterly waves generated by the trade winds there are ocean swells from all directions, but notably from the north in the winter and the south in the summer months. Tide ranges are small, Honolulu having a mean spring tide range of 0.6–0.8 m. Geologically the islands are a series of volcanoes. The main Hawaiian Islands are great shield volcanoes built by successive flows of pahoehoe and aa basalt lavas. Some have had prolonged phases of erosion, followed by renewed eruptions. Hawaiian beaches include creamy white calcareous sand, derived from the tests of micro-organisms, weathered coral, calcareous marine algae, lithic fragments

⊡⊡ Fig. 1.16.1 The Hawaii archipelago consists of eight main islands and numerous smaller volcanic and carbonate islets.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_1.16, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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(typically of Pleistocene skeletal limestone), mollusc and echinoderm fragments (Harney et al. 2000), and black and green sand derived from volcanic material. Ferromagnesian olivine and other basaltic minerals are relatively unstable in the tropical climate of Hawaii, and are reduced by weathering. Calcareous beaches are dominant on all the older Hawaiian Islands where significant coral reef communities have been able to develop. The vast majority of Hawaii’s reefs are of the fringing variety (> Fig. 1.16.2). When an island is young and still volcanically active, lava entering the ocean prevents reef accretion. But moving up the west side of the Big Island where the seafloor is not swept by high swell, the beginnings of fringing reefs attached to the land can be seen and beautiful coral gardens are found. North of the Big Island on the Kihei coast of Maui and extending out from the north Maui shoreline are broad fringing reefs, demonstrating that reef development has a firm foothold on stabilised volcanic coasts. Fringing reefs generally grow in size and become commonplace among the islands north of the Big Island. But among the northwest Hawaiian Islands fringing reefs give way to submerged pinnacles, drowned platforms, and atolls as the volcanic shield structure subsides beneath the waves and the reefs struggle to stay near the surface. Two organisms serve as principal architects of Hawaiian reefs: scleractinian (stony or hard) corals, and coralline and

calcareous algae. There are over fifty species of coral found in the Hawaiian Islands but only a few are common. These grow in a range of forms designed to generally maximize the collection of sunlight and food, and minimize their exposure to stresses made by large waves. Stout branching, delicate branching, platy, encrusting, doming, mounding, these and other terms describe the numerous growth forms assumed by corals as they make the most of their environment. The more abundant Hawaiian genera include rice corals (Montipora species), lobe and finger corals (Porites species), cauliflower or moosehorn corals (Pocillopora species), and false brain corals (Pavona species). Coralline algae and calcareous algae are members of a marine plant group on the reef that deposits calcium carbonate in its tissue. When the algae dies, it leaves a fossil skeleton behind that is hard, whitish, and essentially the same chemistry as the coral. A few species of calcareous algae, such as the Halimeda, are especially abundant in Hawaii and important reef components. Hard plant debris builds up as piles of sediment in reef environments and are important sources of beach sand, making up over half the grains on many Hawaiian beaches. The coralline algae look like coral and grow in a binding and encrusting form on the reef, competing for space with corals. Most coralline algae are red, but there are some exceptions. A visit to any intertidal rocky coast in Hawaii will reveal the ⊡⊡ Fig. 1.16.2 Fringing reefs dominate coastal geomorphology and sedimentary processes on shores not influenced by active volcanism (Larsens Beach, Kauai). (Courtesy Coastal Geology Group, University of Hawaii.)

Hawaii

encrusting coralline community coloring the rocks a brilliant hue in between the rise and fall of the waves. The coastal plains of most Hawaiian Islands hold major calcareous aeolian, littoral, and marine sand deposits formed during and following late Holocene high sea levels  and persistent recent aeolian deposition under seasonal  winds (Fletcher and Jones 1996; Grossman et al. 1998). Sand is also stored on the reef flat in shore-normal reef channels and shallow Pleistocene karst depressions (Fletcher et al. 2008). Longshore transport dominates ­sediment movement on the coast in distinct littoral cells. The central Pacific location of the Hawaiian Islands exposes them to wind and ocean swells from all directions. Sectors of coastline may have rain, wind and wave shadows, and are either protected from, or vulnerable to, wind or wave impact. The four dominant regimes responsible for large waves in Hawaii are: north Pacific swell, trade wind swell, south swell, and Kona storms. The regions of influence of these regimes, outlined by Moberly and Chamberlain (1964), are shown in (> Fig. 1.16.3); a wave rose depicting annual swell heights and directions has been added to their original graphic (Vitousek and Fletcher 2008). Inter-annual and decadal cycles including El Niño Southern Oscillation (ENSO) occurring about every three to four years, and Pacific Decadal Oscillation (PDO) ⊡⊡ Fig. 1.16.3 Hawaii dominant swell regimes after Moberly and Chamberlain (1964), and wave monitoring buoy locations.

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occurring around 20–30 years, influence the ­variability of the Hawaiian wave climate. These large-scale oceanic and atmospheric phenomena are thought to control the magnitude and frequency of extreme swell events. For example, times of strong ENSO may result in larger and more frequent swell. Understanding the magnitude and frequency of extreme wave events is important as they may control processes such as coral development, sediment supply, and beach morphology. In the winter, Hawaii receives large ocean swell from extra-tropical storms that track predominantly eastward from origins in the northwest Pacific. These storms produce waves that travel for thousands of kilometres until reaching the shores of Hawaii. North swell have annually recurring maximum deep-water significant wave heights of 7.7 m with peak periods of 14–18 s. However, the size and number of swell events each year is highly variable – varying by a factor of 2. The annual maximum wave height ranges from about 6.8 m (in 1994, 1997, 2001) to 12.3 m (1988). Occurring about 75% of the year, the trade winds arrive from the east and northeast with an average speed of 25 km/hr and direction 73°. In winter months, the north Pacific high generating these winds flattens and moves closer to the islands decreasing trade wind persistence. Although the number of windy days in summer months

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increases, the mean trade-wind speed in summer and winter months remains similar. The trades generate choppy seas with average wave heights of 2 m (1σ = 0.5 m) and peak periods of 9 s (1σ = 2.5 s) from the northeast. Although these represent nominal conditions, trade-wind swell can exceed 5 m in height and have periods of 15–20 s. Southern swell arriving in Hawaii is typically generated farther away than north Pacific swell. These are usually produced by storms south of the equator near Australia, New Zealand and as far as the Southern Ocean and propagate to Hawaii with little attenuation outside the storm-generated region. South swell occur in summer months (southern hemisphere winter) and reach Hawaii with an annual significant wave height of 2.5–3 m and peak periods of 14–22 s, which are slightly longer than north Pacific swell. Hawaii has occasionally received tsunami generated by circum-Pacific earthquakes, as during the Alaskan earthquake of 1946. Kona storms are low-pressure areas (cyclones) of subtropical origin that usually develop northwest of Hawaii in winter and move slowly eastward, accompanied by southerly winds from whose direction the storm derives its name, and by the clouds and rain that have made these storms synonymous with bad weather in Hawaii. Strong Kona storms generate wave heights of 3–4 m and periods of 8–11 s, along with wind and rain, and can cause extensive damage to south and west facing shores. While minor Kona storms occur practically every year in Hawaii, major Kona storms producing strong winds, large wave heights and resulting shoreline change tend to occur every 5–10 years during the 20–30 year negative PDO cycle. Consequently, positive (warm) PDO, and El Niño phases tend to suppress Kona storm activity. Local relative sea level in Hawaii is not only dependent on global eustatic trends (about 3 mm/yr), but is also affected by subsidence of oceanic lithosphere, which responds elastically to volcanic loading over the hot spot. It is estimated that half of the upward building of Hawaiian volcanoes is reduced by subsidence and that most of the volcanoes have subsided 2–4 km since emerging above sea level (Moore 1987). The main Hawaiian Islands span about 5 million years in age. They are at differing distances from the hot spot, and thus at different stages of subsidence, and so have different relative sea levels. These relative differences in sea levels are demonstrated by modern tide gauge rates and support the view that subsidence is active over the hot spot (Moore 1987). Submerged wave cut notches and benches, raised coral reef, subaerial marine terraces, and alluviated river valleys can be found at various elevations and locations around

Hawaii, indicating oscillations of land and sea level (Fletcher et al. 2008). A fossiliferous marine limestone (known as the Waimanalo Formation) 3–6 m above sea level on Oahu is a typical example of a past stand of the sea. Holocene sea level has been influenced both by eustatic postglacial meltwater as well as equatorial oceanic siphoning associated with the changing postglacial geoid (Mitrovica and Milne 2002). These led to a high sea level (about 2 m) about 3,000 years bp followed by a sea level fall. Tide gauges record a sea level rise since 1900 in Hawaii. There is widespread but variable coastal erosion in the Hawaiian Islands in response to human interference with sand availability and the inferred influence of eustatic rise. For instance, researchers at the University of Hawaii measured the historical rate of shoreline change on every beach on the island of Kauai. Their data reveal that 72 percent of the beaches on Kauai are eroding and the average rate of erosion is 0.3 m/yr. On 22 percent of the eroding beaches, the rate of erosion is accelerating. In pristine coastal areas calcareous sand stored on the low-lying coastal plain is released to the beach as sea level rises, allowing a wide sandy shore to be maintained even as the beach migrates landward. However, the threat to coastal property has led to extensive armoring of the coastline. Artificial hardening is a form of coastal land protection that occurs at the expense of the beach, preventing waves from accessing the sandy reservoirs impounded behind the seawalls and revetments. Thus, efforts to mitigate coastal erosion have created a serious sand deficiency on armored beaches leading to widespread beach loss, particularly on the most populated and developed islands. The need to address this issue is acknowledged by the state and local communities, and the hope is that a broadly scoped management plan will keep a balance between the natural coastal morphology and human resource demands. The islands will be considered in sequence from WNW to ESE (which is also their age sequence): Niihau, Kauai, Oahu, Molokai, Lanai, Kahoolawe, Maui and Hawaii (The Big Island).

2.  Niihau Niihau is a small elongated island (29 km × 10 km) that stretches from SW to NE with a 145 km coastline. It lies at the far NW end of the main Hawaiian islands and is the low lying subaerial remains of a shield volcano built about 4.89 million years ago (Clague and Dalrymple 1987). The horseshoe shaped Lehua Island off the north coast is the product of rejuvenated volcanism and has been sculpted by marine erosion. The steep cliff on the south

Hawaii

side of this relatively small island has been notched with caves at the base. The island tapers to low-lying points that border a wide mouthed bay that opens to the north. The coast of Niihau has a variable morphology. Por­ tions are composed of low volcanic cliffs that have been weathered and shaped by marine erosion, boulder beaches, embayed sandy beach systems backed by 30 m high dunes, erosional sandy coasts where massive slabs of beach rock have been excavated by large swell and stacked on the shore, and small littoral cells with sandy coves.

3.  Kauai Kaui is known for the variety of microclimates that exist throughout the island including temperate regions, dry sand dune complexes and lush river valleys. Volcanic rocks form cliffs and the backshores of beaches, and there are segments of fringing reef cut by paleostream channels, coarse calcareous sandy beaches are separated by rocky points and interspersed with small stretches of boulder coast. The north coast beaches receive large winter surf that may create strong currents. Waterfalls cascade down deep­ ­ly eroded valleys, feeding streams that flow across the beach into the ocean. On the northwest coast one finds the steeply dipping knife-edge ridges and deep erosional V-shaped valleys that have been truncated into spectacular cliffs. The windward and south coasts are characterised by fringing reefs and inter mittent rocky headlands separating distinct littoral cells usually containing calcareous sand beaches. Persistent trade winds and the waves they generate, summer swell from the south, and winter swell from the north all refract around various portions of this coast and influence sedimentary processes. For instance, the Mana Plain on the west shore is a broad accretionary strand plain that originated with falling late Holocene sea level, and converging longshore sediment transport where north swell (refracting from the northwest) and trade wind waves (arriving from the east) deliver sand from two directions at various times of the year.

4.  Oahu Oahu has 180 km of irregular coastline that was greatly influenced by massive landslides that removed about a third of the eastern side of the island and about half the western side. The shape of the island is related to two mountain ranges, the eroded remnants of separate shield volcanoes truncated by the landslides.

1.16

The north shore has massive winter surf (waves 15 m high and more), long sandy beaches, rocky points, and patches of beach rock exposed particularly in winter. The south coast of Oahu has been modified by urban and industrial development (> Fig. 1.16.4). Pearl Harbour formed as the island subsided towards the end of the main shield ­volcano building phase, drowning the river valleys that drain central Oahu. It contains almost 50 km of coastline backed by extensive wetlands through which highly sedimented waters enter the harbour. Along the shoreline, outcrops of 125,000 year-old skeletal limestone reveal fossil corals (> Fig. 1.16.5) in growth position and other paleoreef features that accreted under the higher seas and warmer temperatures of the last interglacial period.

5.  Molokai The island of Molokai has a 142 km coastline, and was formed by at least three shield volcanoes (MacDonald et  al. 1986). Maunaloa (420 m), the West Molokai volcano, is a dry flat-topped shield volcano partly protected from east and NE trade winds by East Molokai (1,514 m), which is wetter and has been cut into deep spectacular valleys and high steep ridges that are lush and vibrantly green. Most of the north coast is cliffed, with prominent ridges and valleys that have been carved by stream ­erosion. There are sectors of vertical cliff cut in volcanic formations.

6.  Lanai Lanai is a single shield volcano that formed from summit eruptions and along three rift zones between 1.2 and 1.46 million years ago, a classic example of a Hawaiian shield volcano with a gently sloping profile. The small sub-circular island has a 76 km coastline and a dry ­climate with minimal stream activity. As on Molokai overgrazing of domestic and feral animals in the nineteenth century and widespread deforestation have drastically changed the stability of the soil. The vegetation has never fully recovered and there is con­siderable wind erosion on the island (Macdonald et al. 1986). Along the northern coast windy conditions have led to the development of a series of low sand dunes behind beaches that are fronted offshore by a narrow fringing reef. The beaches, known collectively as Shipwreck Beach are composed mostly of calcareous sand punctuated by ­outcrops of lithified beach rock. Narrow Polihua Beach has a fringing reef.

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⊡⊡ Fig. 1.16.4 The south coast of Oahu is characterised by intense urban development. World famous Waikiki Beach is shown here in a view taken from the summit of Diamond Head, an extinct volcano on the coast southeast of Waikiki.

⊡⊡ Fig. 1.16.5 The rock shores of O’ahu (photo near Ka’ena Pt.) preserve a record of high sea levels and warmer seas from the last interglacial period (Eemian Interglacial). The yellow line highlights a large coral head that is about 125,000 years old. (Courtesy M. Dyer.)

7.  Kahoolawe Kahoolawe is the smallest of the main Hawaiian Islands, with just 47 km of coastline around a single shield volcano

that formed 1.03 million years ago. It has had a ­complicated history of human exploitation that has affected its terrain and environment, leaving it dry and barren. In the nine­­ teenth and early twentieth centuries, tens of ­thousands of

Hawaii

sheep and goats were raised on the island. Consequent grazing removed most of the vegetation, leaving the soil exposed to high winds, which eventually removed much of it. The coast has lava headlands separated by sandy beaches supplied with sediment from fringing reefs, with pocket beaches of detrital volcanic sand that lie at the mouths of stream gulches that descend from the centre of the island.

8.  Maui The island of Maui is composed of two large volcanoes separated by a low-lying isthmus. The West Maui Volcano (formed 1.6 million years ago) at an elevation of 1,764 m lies to the west of the massive Haleakala ­volcano (formed 0.8 million years ago) with a sum­­mit that reaches 3,055 m. Maui has 193 km of coastline that fringes the two main shield volcanoes and the isthmus. Again the north coast has large ocean waves, and cliffed headlands of volcanic rock alternate with sandy beaches.

9.  Hawaii The island of Hawaii lies over or just north of the Hawaiian hot spot and is composed of five volcanoes and one active seamount: Kohala, Hualalai, Mauna Loa, Kilauea, Mauna Kea and Loihi (located offshore). Of these, only Mauna Loa, Kilauea and Loihi are considered active, while Haulalai is dormant with its most recent eruption ending around 1800–1801.

1.16

The island has 428 km of coastline and is so large relative to the other Hawaiian islands that it is known locally and abroad as the Big Island. Well-developed black and green sand beaches signify the reworking by waves and currents of recent lavas. These beaches are relatively limited along the rough volcanic coastline, and white calcareous sands are restricted because of poor coral reef development due to recent volcanic activity. There is lush vegetation in the NE, where the annual rainfall is 1,500–4,000 mm. The west coast of the Big Island has narrow white sand beaches and lush coral growth at several locations. Locals and visitors alike find that the shoreline has numerous spectacular examples of Hawaiian coastal environments. Kilauea is a large active volcano on the SE flanks of Mauna Loa. On 12 May 2002 the Mothers Day Flow commenced on the south flank of Kilauea and travelled as molten lava through pre-existing lava tubes down to the coast, where it formed a broad lava lobe. When the lava flows into the sea the molten rock (which may exceed 1,100°C) creates steam plumes When pāhoehoe lava enters the ocean for extended periods of time, new land is created in the form of a fan-shaped platform known as a lava bench. Lava pouring into the ocean from either surface flows or lava tubes cools rapidly, usually shattering in the cold water into sand- to block-size fragments. These accumulate along the shore to form a loose foundation that can support overlying lava flows which build a bench above sea level. However, the bench is deceptive; while it looks like solid land, it will in fact eventually collapse into the sea (> Fig. 1.16.6) sweeping unwary visitors to their death. National Park Service personnel usually rope these

⊡⊡ Fig. 1.16.6 A lava bench at East Lae’apuki on the south shore of the Big Island (left) collapsed during the night of 13 August 2007 (right). (Courtesy T. Orr.)

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­ angerous areas off to protect the public. Nonetheless, d foolhardy hikers are seen on these dangerous features, and have been lost when they visited the wrong place at the wrong time.

References Clague DA, Dalrymple GB (1987) The Hawaiian-Emperor volcanic chain: Part 1. Geologic evolution. U.S. Geological Survey Pro­fessional Paper 1350, 5–54 Fletcher CH, Bochicchio C, Conger CL, Engels M, Feirstein EJ, Gross­ man EE, Grigg R, Harney JN, Rooney JJ, Sherman CE, Vitousek S, Rubin K, Murray-Wallace CV (2008) Geology of Hawaii reefs, Chap. 11. In: Coral reefs of the U.S.A. Springer, Berlin, p 435–488 Fletcher CH, Jones AT (1996) Sea-level high stand recorded in Holocene shoreline deposits on Oahu, Hawaii. J Sediment Res 66 (3):632–641

Grossman EE, Fletcher CH, Richmond BM (1998) The Holocene sealevel high stand in the equatorial Pacific: analysis of the insular paleo sea-level database. Coral Reefs 17:309–327 Harney JN, Grossman EE, Richmond BM, Fletcher CH (2000) Age and composition of carbonate shoreface sediments, Kailua Bay, Oahu, Hawaii. Coral Reefs 19:141–154 Macdonald GA, Abbott AT, Peterson FL (1986) Volcanoes in the Sea: the  Geology of Hawaii, 2nd edn. University of Hawaii Press, Honolulu, HI Mitrovica JX, Milne GA (2002) On the origin of late Holocene sea-level high stands within equatorial ocean basins. Quat Sci Rev 21 (20–22):2179–2190 Moberly R, Chamberlain T (1964) Hawaiian Beach Systems. Hawaii Institute of Geophysics Moore JG (1987) Subsidence of the Hawaiian ridge. In: Decker RW, Wright TL, Stauffer PH (eds) Volcanism in Hawaii, US Geological Survey Professional Paper 1350:85–100 Vitousek S, Fletcher CH (2008) Maximum annually recurring wave heights in Hawaii. Pac Sci 62:4:541–553

2.0  Canada – Editorial Introduction

The major geomorphological divisions of Canada are the  Mountain Ranges in British Columbia, the Prairie Provinces, the low-lying Canadian Shield, the Arctic islands and the Maritime Provinces in the east. The Rocky Mountains consist of strongly folded Cainozoic formations, the Prairie Provinces gently-dipping strata, the Canadian Shield Pre-Cambrian crystalline and sedimentary rocks, the Arctic islands a mixture of granitic gneiss mountains and sedimentary lowlands and the Maritime Provinces folded and faulted Palaeozoic and Mesozoic rocks, an extension of the northern Appalachians. Canada has a very long and diverse coastline (McCann 1980; Trenhaile 1990). Its diversity reflects not only the vast size of the country and its geological complexity, but also its range of climate and tides. It has 243,935 km of coastline fronting on the Pacific, Atlantic and Arctic Oceans, of which roughly 172,950 km or 70.9% lies within the Arctic. Pleistocene glaciation was extensive in Canada. The mountains of British Columbia were strongly glaciated, there are residual mountain glaciers, and the coast has numerous fiords. Glacial drift is widespread in the Prairie Provinces, around the Great Lakes and in the Maritime Provinces, and the Canadian Shield includes large areas of ice-scoured rock. The Arctic islands have residual glaciers, notably on Ellesmere, Devon and Baffin Island. Much of Canada has winter snow cover, permafrost is widespread in northern regions, periglacial processes are active in mountain and Arctic regions, and in sea and shore, ice occur at least in winter, along much of the northern and eastern coastline.

1.  Pacific Coast The effects of glaciation and geological structure are evident along the steep coasts of > British Columbia, and in the many fiords. Active cliffs are of limited extent, many steep coast sectors plunging into deep nearshore water. Beaches are also of limited extent, and dunes very rare. South of Vancouver is the large Fraser River delta, forming a lowland that contrasts with the generally steep coast.

2.  Atlantic Coast > New

Brunswick, Nova Scotia and Cape Breton Island consist mainly of gently folded Mesozoic sandstones and shales interspersed with granite intrusions. The New Brunswick coast runs into the Bay of Fundy, and tide ranges increase rapidly to more than 15 m. Wave energy diminishes, and the intertidal zone widens, and becomes muddy and marshy in the Minas Basin and Chignecto Bay  to the north. The southern (Atlantic) coast of Nova Scotia is more exposed. Shore ice is present in winter, but storms generate high wave energy and the tide range is  small ( Gulf of St Lawrence is a large marine inlet that narrows into the estuary of the St. Lawrence River. Wave action is generally moderate, with occasional storms. Shore ice is present in winter, and the tide range in­­ creases from about 2 m around the Gulf to more than 5 m in the St. Lawrence estuary. Prince Edward Island and then east coast of New Brunswick are fringed by sandy barriers fronting lagoons and marshes, the sand having come from glacial drift eroded in cliffs, or reworked on the sea floor. West of the cliffy Gaspe peninsula, the shore is marshy and boulder-strewn, while the northern coast bordering Labrador, is rocky with numerous small valleymouth rias. > St Pierre et Miquelon lie south of the large island of > Newfoundland, which constricts the eastern end of the Gulf. To the west are the Magdalen Islands, ­consisting of sandy ridges and dunes derived from the sea floor.

3.  Arctic Coast > Northern

Canada is defined broadly to include the Labrador coast is similar, with shore ice in winter. Deep

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fiords trench the Pre-Cambrian shield rocks north from Battle Harbour. There are frequent storms, when large waves break on rocky shores, and the tide range is relatively small (Fig.  2.1.3). The strait widens southward towards the U.S. border. The east coast of Vancouver Island has a narrow coastal plain, below 600 m in elevation, less than 30 km

in width, and extending from Johnstone Strait in the north to Victoria in the south (Nanaimo Lowland). The coastline north of Nanaimo consists largely of low-­ gradient broad sand and gravel beaches, derived mainly from erosion of Pleistocene sediment underlying the lowland. More bedrock and mixed sediment-rock intertidal zones occur south of Nanaimo; these are interspersed with local broad beaches. Differential erosion of sedimentary rocks has produced highly indented

British Columbia

2.1

⊡⊡ Fig. 2.1.3 Savary Island consists largely of outwash sands deposited in front of a glacier advancing down the Strait of Georgia about 30,000 years ago.

⊡⊡ Fig. 2.1.4 Goose Spit at Comox on the east coast of Vancouver Island. (Courtesy Geostudies.)

c­ oastlines on southern Vancouver Island and on adjacent islands at the south end of the Strait of Georgia. Only one fiord (Saanich Inlet) occurs on the east coast of Vancouver Island. Abundant sediment is supplied to the littoral zone along the east coast of Vancouver Island, because of erosion of Pleistocene sediment that underlie much of the adjacent lowland. In many areas, these sediment are actively transported along the coast by strong littoral currents and as a result, broad beaches, spits and bars are

common (>Fig. 2.1.4). Rising sea level is expected to increase erosion and sediment transport in these areas (Clague 1989; Shaw et al. 1998). The west coast of Vancouver Island has numerous fiords, many of which extend more than 50 km towards the interior of the island (>Fig.  2.1.5). Much of the shoreline is rugged and rocky, both on the exposed Pacific coast and along the fiords and inlets. Small deltas, marshes and gravelly beaches occur within the fiords. A narrow coastal ­lowland, the Estevan Coastal Plain, extends more than 270 km along

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British Columbia

⊡⊡ Fig. 2.1.5 Fiord on the west coast of Vancouver Island.

⊡⊡ Fig. 2.1.6 Long Beach, western Vancouver Island.

the west coast of Vancouver Island from the mouth of Juan de Fuca Strait to Brooks Peninsula in the north. The only extensive beaches on the west coast of Vancouver Island occur within the Estevan Coastal Plain (>Fig.  2.1.6). Most of the beaches have developed in areas underlain by thick Pleistocene sediment where offshore gradients are low. At Florencia and Wickaninnish bays,

between the towns of Ucluelet and Tofino, wide sandy beaches have developed through erosion of uncon­solidated sediment. Florencia Bay is a crescent-shaped embayment with a small cuspate foreland in the middle and rocky headlands at its northwest and southeast ends. The beach, about 5 km in length, is backed by a steep cliff of Pleistocene sediment.

British Columbia

2.1

⊡⊡ Fig. 2.1.7 Cliff in Pleistocene sediment, Point Roberts south of Vancouver.

6.  Fraser Delta

References

Much coastal geoscience research in British Columbia has been directed to understanding the Fraser River delt south of Vancouver (Clague et al. 1998; Barrie and Currie 2000; Kostaschuk 2002; Mosher and Thomson 2002). The active, western front of the delta is about 23 km long and is fed by a network of distributary channels. Perhaps the most conspicuous features of the delta are sand and mudflats up to 6 km wide, locally crossed by distributary channels and bordered seaward by a foreslope with an average gradient of 1.5°, but with slopes up to 23° along its inner, shallower portions. Large amounts of Fraser River sediment are carried to the river mouth during the summer freshet. The western delta front is separated from the inactive southern part of the delta by Point Roberts Peninsula, a ridge cored by Pleistocene sediment (>Fig.  2.1.7). The upper foreslope of the active delta is primarily muddy north of the main distributary channel and sandy to the south, reflecting the dominant ­northward transport of Fraser River suspended sediment during floods.

Anderson JL, Walker IJ (2006) Airflow and sand transport variations within a backshore-parabolic dune plain complex; NE Graham Island, British Columbia, Canada. Geomorphology 77:17–34 Barrie JV, Currie RG (2000) Human impact on the sedimentary regime of the Fraser River delta, Canada. J Coastal Res 16:747–755 Clague JJ (1989) Sea levels on Canada’s Pacific coast: past and future trends. Episodes 12:29–33 Clague JJ, Bornhold BD (1980) Morphology and littoral processes of the Pacific coast of Canada. In: McCann SB (ed) The coastline of Canada. Geol surv Canada Paper 80–10. pp 339–380 Clague JJ, Luternauer JL, Mosher DC (1998) Geology and natural hazards of the Fraser River delta. Geol Surv Canada Bull 525:270 Holland SS (1964) Landforms of British Columbia, a physiographic outline. Br Columbia Dep Mines Petrol Resour Bull 48:138 Kostaschuk R (2002) Flow and sediment dynamics in migrating salinity intrusions, Fraser River estuary, Canada. Estuaries 25:197–203 Mosher DC, Thomson RE (2002) The Foreslope Hills; large-scale, finegrained sediment waves in the Strait of Georgia, British Columbia. Marine Geol 192:275–295 Shaw J, Taylor RB, Forbes DL, Ruz MH, Solomon S (1998) Sensitivity of the coasts of Canada to sea-level rise. Geol Surv Canada Bull 505:79

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2.2 New Brunswick and Nova Scotia

Eric Bird

1.  Introduction New Brunswick and Nova Scotia are geologically part of the northern section of the Appalachian system, and consist predominantly of deformed Palaeozoic sedimentary rocks trending SW–NE. The region was strongly glaciated in Pleistocene times, notably in the Wisconsin phase. Glaciated valleys and morainic deposits are extensive in lowland areas. The climate is cold in the winter months, when shore ice may form in sheltered bays (Owens and Bowen 1977). Atlantic ocean swell, attenuated across a broad con­ tinental shelf, reaches the coast of southeastern Nova Scotia, and is refracted into the Bay of Fundy and the New Brunswick coast. There is exposure to frequent storm waves generated during the passage of cyclonic storms from west to east across the northern Atlantic. Mean spring tide ranges are commonly less than 2 m on the Atlantic coast of Nova Scotia, but they are larger in the Bay of Fundy.

2.  New Brunswick From the Maine border, where Croix River flows into Oak Bay, the New Brunswick coastline is irregular, with bays and inlets between rocky headlands, as at Point Lepreau. Much of the coast is cut across crystalline rocks, forming a hilly hinterland, but there is a sector of Carboniferous rocks at Quaco. The coastline borders the Bay of Fundy, and exposure to Atlantic Ocean waves diminishes northeastward. At Saint John the mean spring tide range is 7.4 m, and it increases along the coast towards the upper branches of the Bay of Fundy, exceeding 10 m in Chignecto Bay. Coasts with tide ranges exceeding 6 m are classed as megatidal. Stacks cut in Carboniferous conglomerates in Blacks Harbour show unusually deep and tall abrasion notches due to the large tide range (>Fig.  2.2.1). Tidal bores run upstream in the Petitcodiac River, near Moncton at the head of Chignecto Bay. The tide is semidiurnal, with only a small diurnal inequality, and tide range has increased over the last 6,000 years as the dimensions of the bay have changed because of eustatic sea level rise and ­isostatic

recovery. Much of the interest in the tides and their influence on sediment dynamics is related to the possibilities of tidal power generation. The megatidal Cumberland Basin has wide muddy shores exposed at low tide and extensive salt marshes. There has been much land reclamation. A submerged forest at the mouth of the Missaguash River, near Fort Lawrence, is exposed at low spring tides and submerged by up to 10 m of sea water at high spring tides.

3.  Nova Scotia The coastline of Nova Scotia includes the megatidal Bay of Fundy, with maximum tide ranges rising to more than 16 m at the head of the bay, and the Atlantic coast which is generally low lying, mesotidal and dominated by occasional storm waves. Geologically Nova Scotia consists of two tectonic plate segments, the Avalon in the north and the Meguma in the south, separated by the major Cobequid-Chedabucto fault system which runs through Parrsboro and Five Islands ESE to Chedabucto Bay. South of this fault system, an Early Jurassic extrusion of basalt (the North Mountain Formation) forms headlands at Cape D’Or, Partridge Island, Wasson Bluff and Five Islands, and the long ridge of North Mountain and Digby Neck south of the Bay of Fundy. During the Quaternary there were several phases of glaciation, when ice moved from north to south across Nova Scotia. River valleys were deepened into troughs, and when the glaciers melted they left extensive deposits including moraines, drumlins (ice-moulded drift mounds) kames (gravelly, low hummocky mounds deposited on or within the ice), eskers from subglacial streams and glacifluvial outwash plains. The last (Wisconsin) glacial phase ended about 16,000 years ago when the climate became warmer and sea level rose in the Late Quaternary marine transgression. Spring tide range in the Bay of Fundy increases from 5.75 m at Brier Island to 13–14 m in Chignecto Bay and Amherst Harbour has a range of 11.6 m. At Burntcoat

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⊡⊡ Fig. 2.2.1 Stacks on the New Brunswick shore of the Bay of Fundy, showing unusually tall abrasion notches on a megatidal coast. (Courtesy Gudrun Reichert.)

Head on the south coast of the Minas Basin spring tide range is 16.3 m. These are among the highest tide ranges in the world, possibly equalled or exceeded in Ungava Bay, Labrador. Sea level changes in Nova Scotia have been influenced by glacio-isostatic movements of the land. The Late Quaternary marine transgression was rapid between 8,000 and 2,000 years bp, and then decelerated. Tide gauge records show a continuing rise of about 3.2 mm/ year at Halifax, 4.7 mm/year at Yarmouth and 3.8 mm/ year at North Sydney on Cape Breton Island. Rising sea level in the mid-Holocene between 7,000 and 4,000 years bp amplified the tide range in the Bay of Fundy. Coastal processes are modified in winter by the formation of sea ice, particularly in the Bay of Fundy, where it begins to form in December, attains a maximum in February, and melts in March. Shore ice persists in estuaries into April. Sea ice, often containing boulders, gravel and sand, and darkened with finer sediment, is agitated by wave action and contributes to the erosion of cliffs and shore platforms. It is also rafted on to intertidal flats and salt marshes, contributing sediment as it melts. Alter­natively, sediment trapped in sea ice may be carried away in ice blocks floating offshore, and deposited on the sea floor. The boundary between New Brunswick and Nova Scotia follows the Missaguash River, which flows into the head of the Cumberland Basin. To the south is Lusby Marsh, dissected by meandering tidal creeks, on a coast where the spring tide range is more than 15 m. The wide estuary of the River Hebert comes in from the south, joined by the Maccan River estuary, which is deflected southward

by Amherst Point, a spit that widens southward, backed by salt marshes. Sedimentation has maintained these in a period of rising sea level (Shaw and Ceman 1999). Extensive salt marsh areas around the upper Bay of Fundy have been embanked (dyked) and reclaimed as pastureland. This megatidal coast has very wide shores exposed at low tide, with a tide-dominated morphology of sand bars and ripples of varying dimensions. There is a notable submerged forest near Fort Lawrence, in the Missaguash River estuary, exposed at low spring tides and submerged by up to 10 m of sea water at high spring tides. The megatidal environment results in tidal bores in the Missaguash River and the River Hebert at the head of the Cumberland Basin. Much of the interest in the tides and their influence on sediment dynamics is related to the possibilities for tidal power generation. Between the Cumberland Basin and these estuaries is the broad anticlinal Minudie promontory, marshy at its northern end and becoming hilly southward. At Downing Head the coast becomes cliffy, with exposures of Late Carboniferous (Cumberland Group) rocks that extend southward past Joggins and out to Ragged Point. The cliffs at Joggins, between Hardscrabble Point (Coal Mine Point) to the north and Ragged Reef Point to the south are known internationally as a classic coastal geology site -the Joggins Fossil Cliffs Provincial Park (> Figs. 2.2.2 and > 2.2.3). The vertical cliffs and wide intertidal shore platforms are cut in southward dipping grey Carboniferous sandstones and shales, with occasional coal seams and limestones. The strata exposed in the cliff face show the southward dip, with some faulting, on the flank of the

New Brunswick and Nova Scotia

Minudie Anticline, declining as it descends into the Atholl Syncline to the south. The basal Boss Point Formation outcrops in the cliffs and shore platform north of Little Cove and passes southward beneath the Joggins Formation, which in turn dips under the Springill Mines and Ragged Reef Formations to the south. The shore is accessible at Lower Cove, or down steps into a chine at the mouth of Bell Brook, where two

⊡⊡ Fig. 2.2.2 Gravel-strewn shore platform at Joggins, fronting a cliff in southward-dipping Carboniferous sandstone. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.3 Cliffs near Hardscramble Point. Joggins. (Courtesy Geostudies.)

2.2

streams descend dip planes and converge. At Bell Brook there is a shingle upper beach with grit and sand, some coal pebbles and bright red bricky sandstone fragments sorted into shore-parallel zones. The cliffed coastline cuts across the geological strike, and shows variations related to harder and softer outcrops. The intertidal shore platform shows similar variations, and 2.5° to 3°, spring tide range here being 13 m. The platform is strewn with

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gravel, including coal fragments, particularly north of the piles and timbers which are the remains of the dock at Joggins where coal was loaded. The lower shore is sandy. Ribs of hard sandstone run out across the shore and form minor cuestas, with steep escarpments facing north and gentler dip-slopes southward between lower and flatter zones on softer intervening outcrops. The shore platform has been cut by wave abrasion as the tide rises and falls by the action of driven sea ice action, and frost weathering. Near Hardscramble Point the Upper Carboniferous rocks are overlain by glacial drift. Coal has been mined here, and the coal-bearing sediment include upright casts of tree trunks perpendicular to the dipping strata. The rock formations are fossiliferous, yielding grey coral fossils, crinoid stems, shells and leaf impressions. In 1842 Charles Lyell visited here, and with local experts John William Dawson and Abraham Gesner debated the origin of coal. Dawson and Lyell found vertebrate bones in a lithified fossil tree stump. It has been suggested that the Joggins Fossil Cliffs should become a UNESCO World Heritage Site because of their geological significance. The Carboniferous rocks are capped in several sectors by unconsolidated Pleistocene glacial drift deposits. The cliff section near MacCarrons Brook shows three stages of glacial deposition – an upper yellow Shulie Lake till over middle grey Joggins Till resting upon lower reddish muddy McCarron Brook Till.

The generally cliffed coast continues SW to Cape Capstan, beside Apple River Bay, where The Bar is a midbay barrier spit fronting grassy salt marsh in the estuary of Apple River. The spit passes south into a beach with cliffs behind curving Edgells Beach, which has shore plat­forms cut in grey Carboniferous sandstone. Beyond Pudsey Point the coast rises and steepens southward to the bold headland at Cape Chignecto. At Cape Chignecto the coast turns ENE and becomes very steep, dissected by several deeply-incised river valleys. At McGahey Brook it curves ESE and begins to decline and at West Advocate the sloping cliffs (>Fig.  2.2.4) descend to a shingle beach piled with driftwood at high tide and interrupted by outcropping ledges of pink Triassic sandstone forming a sloping abrasion platform. The beach continues as a barrier spit fronting a former high tide salt marsh enclosed by a dyke and reclaimed as meadowland - then Advocate Harbour – which has a narrow outlet between paired spits. The hinterland is hilly and wooded on the Carboniferous formations. South of Advocate Harbour the land rises to Cape D’Or, a bold cliffy headland of Jurassic basalt, with high partly overhanging cliffs on the western side. This is the Cape d’Or Scenic Area. The headland takes its name from the cliffs stained yellow by copper (not gold) ores. A steep grassy bluff descends to a low promontory with a projecting basalt ridge and southern cliffs with ledges cut across columnar basalts dipping south. The large tides produce a pronounced eddy, the ebb current swinging round to run

⊡⊡ Fig. 2.2.4 Cliffs in Triassic sandstone at West Advocate. (Courtesy Geostudies.)

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westward inshore and forming waterfalls over basalt reefs. A similar eddy develops on the rising tide on the western side, and the two meet in the rough waters of the Dory Rips. On the eastern side of Cape d’Or is Horseshoe Cove, where a shingle beach impounds a stream-mouth pond. Cliffs of southward-tilted basalt run out eastward to Cape Spencer – also of basalt (>Fig. 2.2.5). The high forested basalt ridge of Cape d’Or is backed by a low corridor with a fault along which rock formations have been displaced laterally (much like the Great Glen in Scotland). In Greville Bay, to the east, low bluffs of glacial drift and grey Carboniferous shales are interrupted by the mouths of several incised valleys, some with salt marsh segments at high tide level – as at Wards Brook which drops through a gorge to a marshy tidal inlet. At Port Greville a coastal terrace is bordered by a bluff that descends to a shingle beach. Fox River and a series of small streams are incised into a hilly coastal fringe, and descend to marshy inlets, some with sand and shingle barrier beaches across their mouths, as at Fox Point and the mouth of Diligent River. The coast rises to steeply sloping cliffs at Bull Bluff, and at Black Rock a longshore spit of grey shingle has grown eastward to shelter a high tide lagoon. Cape Sharp is another headland of Jurassic basalt. Black Rock is also the name of a small rocky island offshore, round which strong tidal currents swirl. Barrier spits enclose tidal lagoons (>Fig.  2.2.6) and tombolos link nearshore islands. The Minas Channel narrows to about 5 km here, with views

⊡⊡ Fig. 2.2.5 Basalt cliff on promontory near Cape D’Or. (Courtesy Geostudies.)

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across to North Mountain, a high basalt ridge ending westward at Cape Split. The Late Quaternary marine transgression flooded the land to various levels, declining eastward from at least 35 m near Cape Chignecto to 0 m at Truro, as recorded by proglacial deltas. In the Parrsboro region there are terraces of meltwater outwash and postglacial fluvial sediment dissected as the sea dropped back from the Late Glacial marine limit. The major Cobequid-Chedabuto Fault system runs inland behind Parrsboro, parallel to the coastline and marked by a scarp in the Upper Carboniferous rocks on the hard Horton Group beside the softer downfaulted Parrsboro Formation. The strait between Cape Split and Cape Sharp, west of Parrsboro, leads into the Minas Basin, where mean spring tide range is over 15 m. At Burntcoat Head on the south coast of the Minas Basin spring tide range is 16.3 m. These are among the highest tide ranges in the world, possibly equalled or exceeded in Ungava Bay, Labrador. The macrotidal coast has very wide shores exposed at low tide, with a tide-dominated morphology of sand bars and ripples of varying dimensions. The bordering coast has extensive glacial drift deposits and soft sandstones, locally cut into cliffs which are reached by waves only briefly at high spring tide, and have a sloping profile. There are beaches of sand and shingle, derived from the eroding cliffs (>Fig.  2.2.7), ­passing ­seaward into gravel with a thin mud veneer. Amos and Long (1980) estimated a net annual sediment input into

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the Basin of 4.8 million m3, of which 3.09 million m3 is  derived from cliff recession. In addition to beaches, there are longshore spits, shaped by a combination of  high tide waves and currents. Erratic boulders (>Fig.  2.2.8) have been stranded on the shore, and in sheltered bays there are segments of salt marsh in the upper intertidal zone; these become more widespread in the upper bays and estuaries.

East of Cape Sharp is the curving shore of Parrsboro Roads, backed by high sloping bluffs. These curve out to East Bay, where a broad tombolo of sand and gravel links forested, flat-topped Partridge Island – also of dark basalt – to the mainland. East Bay has a curving beach of fine platy shingle, with overwash fans along the crest, bordering the tombolo, a grey gritty plain that includes a tidal creek and sparse marsh, with an outlet on the western

⊡⊡ Fig. 2.2.6 Barrier spit and tidal lagoon at Black Rock. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.7 Sloping cliff of glacial drift near Sand Point on the north coast of the Minas Basin. (Courtesy Geostudies.)

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⊡⊡ Fig. 2.2.8 Erratic boulders on the shore at Five Islands. (Courtesy Geostudies.)

shore close to the island. Across the Minas Channel is Cape Blomidon, with an upper vertical cliff of columnar Jurassic basalt above a steep slope on red Triassic rocks. Parrsboro has a spring tide range of 11.6 m. Farrells River opens into a branching estuary at Parrsboro, and to the east, low cliffs extend to Clarke Head, and on past Wasson Bluff and McKay Head to the Five Islands Provincial Park. The coast east of Wasson Bluff has extensive glacial drift deposits and soft sandstones, locally cut into cliffs which are reached by waves only briefly at high spring tide, and has a sloping profile. There are beaches of sand and shingle, derived from the eroding cliffs, passing seaward into gravel with a thin mud veneer. As well as beaches there are longshore spits, shaped by a combination of high tide waves and currents. Erratic boulders have been stranded on the shore, and in sheltered bays there are segments of salt marsh in the upper intertidal zone. At Five Islands the high islands form a dissected extension of the ridge running west from Economy Mountain. A cliff south of the lighthouse shows the red sand­stones and mudstones of the early Jurassic McCoy Brook sedimen­ tary Formation faulted against dark grey North Moun­tain basalt to the west. On the northern side of the promontory, the red McCoy Brook Formation is capped by glacial drift and outwash, while on the southern side the upfaulted North Mountain basalt caps sloping red cliffs in the Triassic Blomidon Formation. At the end of the mainland cliff is the Old Wife basalt sea stack. The backshore has sectors of Spartina marsh.

The area around the Five Islands drains out as the tide falls to reveal wide sand flats and large sand bars. The sand bars, which are generally composed of medium to coarse sand, are elongated and may be several kilometres in length, with an amplitude of 10–15 m. Their surfaces are covered with large ripples generated by tidal currents, which reach velocities in excess of 1.5 m/s. It has been suggested that the large sand bars are equilibrium forms subjected to a dispersal system of alternating flood and ebb transport. Tidal currents and sediment transport pathways maintain the pattern of sand bars and channels in Cobequid Bay at the head of the Minas Basin (Knight 1980). Present on the sand bars are four kinds of bedform: current ripples (height about 0.05 m, wave length about 0.3 m), two kinds of megaripples (height 0.05  m, wave length 1.0 m and height 0.05 m, wave length 15 m) and sand waves (height 0.2–3.0 m, wave length 15–200 m) (Dalrymple et al. 1978). Most sand waves have megaripples superimposed on them (>Fig. 2.2.9); and ripples are superimposed on all larger bedforms. As a result of these and similar studies, the Minas Basin is a key reference area for models of bar and ripple formation. The combination of coarse grain size and mobile sediment means that the intertidal fauna of the sand bars is sparse, but the muddier tidal flats support a wide variety of organisms. Resuspension of biodeposits contributes significantly to turbidity of the Minas Basin waters. East of Economy Mountain at Carrs Brook are cliffs and stacks in the Triassic Wolfville Formation. Economy

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⊡⊡ Fig. 2.2.9 Sloping cliff of glacial drift near Sand Point on the north coast of the Minas Basin. (Courtesy Geostudies.)

Point is another wooded peninsula with red earthy cliffs on its western coast. The coast to the east is generally low, with the mouths of several rivers draining down from the Cobequid Mountain. Portapique River is gravelly, and Bass River opens into a marshy bay. At Saints Rest an intertidal deposit includes tree stumps dated 8,180 and more than 4,400 years bp, indicating a lower sea level. At Highland Village low cliffs cut in pink clay (Triassic Wolfville Formation) stand behind wide sand and mudflats exposed at low tide, and the coast is low and marshy east of Spencers Point. Cobequid Bay has water area of 306 sq. km at high tide, and intertidal area of 186 sq. km exposed at low tide. It narrows eastward the combined estuary of North River and Salmon River at Truro, and has extensive intertidal sand bars (Dalrymple et al. 1975). Amos and Long (1980) estimated a net annual sediment input into the Minas Basin of 4.8 million m3, of which 3.09 million m3 was derived from cliff recession. The southern coast of Cobequid Bay is also low-lying, and the funnel-shaped estuary of Shubenacadie River comes in from the south. West from Salter Head the coast is embayed, with many small creeks opening through salt marshes. Cobequid Bay widens westward into Minas Basin, which receives the broad estuary of the Avon River. Cornwallis River drains the north-eastern part of the Annapolis Valley into the Minas Basin. Each of these estuaries has slumping banks of brown clay exposed at low tide, when megaripples are exposed in the intertidal zone.

Rising tides form diurnal tidal bores that flow upstream at mid-tide. During spring tides a tidal bore rushes upstream into the rivers that enter Cobequid Bay – particularly the St Croix River near Windsor, the Shubenacadie and the Salmon River at Truro. The river channels are bordered by high artificial levees. There are several historic river ports, including Hantsport on the River Avon and Port Williams on the Cornwallis River. Wolfville Harbour drains out at low tide. The western coast of Minas Basin has sloping cliffs of red Triassic rock, which passes northward beneath Jurassic basalt forming the bold cliffs of the Blomidon Peninsula, which show columnar structures (>Fig. 2.2.10). This marks the beginning of the long ridge of Jurassic basalt (North Mountain) which runs WSW to Brier Island, with marine gaps at Digby Gut, Petit Passage and Grand Passage. It forms a cuesta with an escarpment facing SSE across the Annapolis Valley, the tidal Annapolis Basin, the Digby isthmus and St Marys Bay, and a dip-slope declining towards the Bay of Fundy coast, where seaward dipping layers of basaltic lava outcrop on the shore. At its NE end it curves sharply round Scots Bay so that the escarpment forms the southern coast of the Minas Channel. It ends at Cape Split, with an outlying stack. There are roaring turbulent currents (up to 4 m/s) here on the middle of a rising tide. From Cape Split the Scots Bay coastline consists of seaward-dipping basalt with coves at the mouths of small valleys drained by short streams and wide intertidal sand and mudflats. The streams lengthen as the coast curves

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⊡⊡ Fig. 2.2.10 Cape Blomidon, showing Long Mountain basalt over Triassic sediment. (Courtesy Geostudies.)

southward, then WSW to the long straight section along the south coast of the Bay of Fundy. The view from the hinterland slope is across the Bay of Fundy to outlying Isle Haute, Cape Chignecto, Cape D’Or and Cape Split to the east. Valleys descend to small bays, several of which have small harbours for fishing boats, sheltered by massive, boulder-armoured breakwaters, as at Halls Harbour. Low cliffs and sloping ramps of basalt with scattered boulders are interrupted by curving bays of well-sorted grey shingle derived from basalt and glacial drift, as at Hampton Beach. Parkers Cove has another harbour with stone breakwaters and massive granite boulders on the eastern side. The Annapolis Valley runs ENE-WSW, and the Annap­ olis River has a meandering estuary that opens into the  tidal Annapolis Basin. At Annapolis Royal the tide range is about 7 m, utilised by a hydroelectric station in a  narrow sector of the Annapolis River. The Annapolis Basin widens towards Digby, and Digby Gut is a steepsided channel through the North Mountain coast range, forming an outlet to the Bay of Fundy (>Fig.  2.2.11). The semidiurnal tide has a range of 8 m at Digby, and a broad shore zone is exposed around the Annapolis Basin at low tide. Basalt layers dip NW on either side of Digby Gut, and sloping shore ledges follow the dip below the lighthouse at Point Prim on the western side. The Bay of Fundy coast continues to the WSW with valley-mouth coves between headlands with ramps of dipping basalt. Broad Cove

c­ ontains a harbour enclosed by breakwaters in a bay with a grey cobble beach curving round to lava cliffs. Gullivers Cove is a bay with an upper beach of cobbles and boulders, mainly basalt and also of other rock types from glacial drift, partly held by a backshore stockade of timber, with a timber groyne. The North Mountain cuesta is set forward along a fault line at Gullivers Head, and the next section is known as Digby Neck. There are low cliffs cut in glacial drift and successive lava ledges along the shore. The outermost of these on Gullivers Head is an almost flat shore platform, draped with seaweed. Goose­ berry Cove has a grey boulder and cobble beach with a high cliff of seaward dipping basalt on the SW side. The cliffs and storm-piled beaches indicate that wave action can be strong on this Bay of Fundy coast, but the large tides reduce effective wave energy. At Centreville a valley occupied by glacial drift contains incised Trout Creek, which flows out through a gravel beach into a tidal harbour, enclosed by tall breakwaters with wooden pilings, stone filling and boulder armouring. The sloping basalt coast continues WSW to the gap at Petit Passage, resumes in Long Island to Grand Passage, and ends on Brier Island. The inner coast, facing St. Marys Bay, has low cliffs and bluffs truncating at the base of the basalt dip-slope, small coves and some beach fringes. At The Sea Wall the coast is lined by granite boulders that have been dumped across the mouth of a salt marsh bay. The coast then curves out to Red Head, where cliffs have been cut in

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red Triassic sandstone, which underlies the Jurassic basalt. At the head of St. Marys Bay the coast of the Digby isthmus is low-lying on glacial drift, and at low tide the northern part of the bay has extensive brown gravelly mud and sand flats. The southern shores of St. Marys Bay are at first lowlying and marshy, with cliffs cut in glacial drift at Gilberts Head. To the WSW, marshes give place to beaches and longshore spits of sand and shingle derived from glacial

drift, as at Brooks Beach. The estuary of Sissiboo River enters below Weymouth. At Belliveau there is a little harbour, and a long curving shingle beach to west becomes a lobate foreland with multiple beach ridges at Pointe a Major (>Fig. 2.2.12). To the south, Rivière Grosses Coques flows out under an old road bridge to a broad salt marsh – Marais de Grosses Coques, behind a shingle barrier spit – the ­estuary opening near the SW end. At Ticken Cove a grassy dune ⊡⊡ Fig. 2.2.11 Digby Gut, a gap in the Long Mountain ridge. (Courtesy Geostudies.)

⊡⊡ Fig. 2.2.12 Gravel beach and parallel beach ridges at Pointe a Major. (Courtesy Geostudies.)

New Brunswick and Nova Scotia

barrier fronts a bay lagoon, and Church Point shelters a salt marsh, with gravel shoals backed by a beach with dunes fronting a lagoon. At Little Brook there are gentle grassy slopes down to the coast, and at Comeauville and Saulnierville low cliffs in glacial drift are fronted by shingle beaches. Southward along the coast, wave energy increases with widening fetch and diminishing tides as the Bay of Fundy opens into the Gulf of Maine and the Atlantic Ocean. Tide range diminishes from 8.5 m at Digby Gut to 5 m at Cape St. Marys. At Smugglers Cove (>Fig. 2.2.13) there is a bold cliffed coast cut in Lower Palaeozoic mudstones with steep to vertical cleavage. Some of the cleavage planes are exposed in the cliff face, disintegrating into angular gravel, which forms a cobble beach. At Cape St Marys the lighthouse stands on cliffs cut in the cleaved mudstone, which now trends at right angles to the coast so that there are narrow promontories between deep clefts and an arête. South of Cape St Marys is a fishing harbour, and a long curving sandy beach backed by dunes in Mavilette Beach Provincial Park. This ends in a cliff cut in brown glacial drift, followed by more cliffs in grey mudstones. There are several lagoons behind grassy barriers of sand and shingle. At Salmon River the stream has a narrow outlet to the sea, bordered by a wall of wooden piling. The coast runs southward with many inlets at the mouths of river valleys, past Yarmouth to Chebogue Point. Tide ranges diminish along the south-west coast Nova Scotia to 3.3 m

⊡⊡ Fig. 2.2.13 Bedding planes exposed in cliff face at Smugglers Cove. (Courtesy Geostudies.)

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at Yarmouth, and wave action becomes more effective as the tides decline. There are cliffs cut in glacial drift and derived beaches and spits of sand and gravel. Long narrow peninsulas, as at Pinkneys Point, separate marine and estuarine inlets and the Tusket Islands are an outlying rocky archipelago. The Atlantic coast of Nova Scotia is indented, with long narrow peninsulas and rocky ridges separating elongated marine and estuarine inlets at valley mouths and many rocky islands. The rock formations are mainly Lower Palaeozoic (Cambrian and Ordovician) metasediment (sandstones and shales with some schist and gneiss), and several granite intrusions. The inlets were river valleys deepened by glaciation (an ice sheet that flowed N-S), then submerged by the Late Quaternary marine transgression. There is a mantle of glacial moraine with some drumlins, locally truncated by sea cliffs. Beaches, spits and barriers of sand and gravel have been derived largely from these glacial drift deposits. Coastal waters are generally ice-free in winter, but ice may form in estuaries and lagoons. Wave action, mainly from the SE, is strong in frequent stormy periods, and there are occasional storm surges of at least 1 m. Tides are semidiurnal, mean tide range diminishing from 3.7 m in the SW to 1.5 m at Halifax and less than a metre in the north of Cape Breton Island. Lahave River, downstream from Bridgewater, opens to the sea beside the Kingsburg peninsula. There are barrier beaches and spits enclosing lagoons, as in Hartling Bay,

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and at Rose Bay a spit shelters a salt marsh on Conrad Island. North of Ovens Point is gravelly Cunards Beach, and there are cliffs cut in slaty mudstones with clefts and Ovens Caves, cut out along cleavage planes. Some of the caves were adits from historical gold mining. St Margaret’s Bay is wide, with many wooded islands, and Peggy’s Cove to the east has sloping granite shores with glacial striations and a prominent lighthouse. At Sambro there are intricate peninsulas and inlets with seaweeddraped bouldery shores and occasional sandy beaches. Crystal Crescent Beach has grey sand and scattered boulders, with granite shore outcrops and patches of grassy marsh. Around Ketch Harbour the inlets are edged with pale grey granite boulders, and Portuguese Cove and Herr­ ing Cove are more inlets fringed by granite boulders and small sandy beaches. In all these locations the granite has glacially-smoothed slopes down into the sea, without shore platforms. Halifax and Dartmouth stand on either side of a broad marine inlet, Halifax Harbour. Upstream, through The Narrows, is the Bedford Basin, a former lake that became submerged during the later stages of the Late Quaternary marine transgression. As on McNab’s Island, erosion of bluffs in glacial drift has occurred, in the marine inlet south of Halifax (Manson 2002). On the outer coast east of Halifax there have been rapid changes. Cliffs are receding, particularly on drumlin headlands and islands, and there has been associated landward migration of beaches, spits and barriers, with overwashing and breaching which results in the movement of sediment into lagoons that have become more saline, with replacement of freshwater swamps by salt marshes. These rather rapid changes have been accelerated by the rising sea level indicated on the Halifax tide gauge (Taylor et al. 1996). Halifax Harbour is a marine inlet. To the east, the peninsula south of Dartmouth has an outer coast with headlands at Hartlen Point and Osborne Head. An intervening bay, Cow Bay, has a gravelly barrier beach (Silver Sands Beach) with multiple cobble beach ridges enclosing a lagoon – Cow Bay Lake – in a former embayment. On the eastern side is a cliffed drumlin, and there are records of former drumlins, now planed off as shoals, which were the main source of sediment for this barrier beach. In 1900 Cow Bay Lake was freshwater, with an outlet to the sea at the western end of the beach. However, after this was closed by storm waves, the lake overflowed and cut a new entrance to the east. A breakwater was then built to provide a boat harbour, and Cow Bay Lake has since been brackish. The barrier beach was reduced in width by sand and gravel extraction

in the 1950s and 1960s, and has migrated landward as the result of storm wave overwash. Cole Harbour is a branched embayment with bouldery shores with patches of grass marsh, backed by woodland that was damaged by Hurricane Juan on 29 September 2003. There is a sandy barrier spit at its mouth (Rainbow Haven) and extensive salt marshes, with some sandy beaches. To the east is Conrad Head, with a sandy cliff, and beyond it Conrad Beach – a sandy barrier beach backed by dunes and the marshes of West Marsh – curving round to cuspate Fox Point. Like Silver Sands Beach this was derived from sediment eroded from cliffed drumlins, and reduced by the quarrying of beach gravel in the mid-twentieth century. The barrier was breached during a storm in 1962 and remained open until 1989–1990, since when dunes have developed across the former inlet. The coastal features can be viewed from the top of Lawrencetown Head, a cliffed drumlin. The cliff has been receding at up to 0.8 m/year. Lawrencetown Beach has low grassy dunes, and the shingle beaches are fronted by sandy shores exposed at low tide. The sediment has been derived from cliffed drumlins and morainic drift deposits. A barrier beach of sand and gravel encloses elongated Law­ rencetown Lake. At Fishermans Reserve there are shacks and boats on a projecting spit, the north shore protected by massive boulders. To the east Rudeys Head is a cliffed drumlin. The barrier-enclosed lagoons are like New England ponds, marshy with Typha along their inner shores. Gaetz Head is another truncated drumlin headland (>Fig. 2.2.14). When these drumlin headlands were cut back, the barrier beaches derived from them have retreated landward. Cape Antrim, a drumlin headland, disappeared as the result of cliff recession at rates that attained more than 12 m/year in the 1970s, and the adjacent barrier beach receded at a similar rate. Chezzetcook Inlet has shoals and spits trailing NW behind islands. It is typical of the wave-dominated salt marsh fringed estuaries with only minor fluvial inflow on the coast east of Halifax (Carter et al. 1992). There are salt marshes with Spartina patens in the upper marsh, and Spartina alterniflora at lower levels, fringed by mudflats exposed at low tide. Some of the sediment has been supplied by rivers, but much has come from marginal shore erosion, soil creep on slopes (increased following forest clearance) and overwash of spits and barriers at the seaward end. Story Head Beach is a low gravel barrier beach 40–60  m wide that extended from Story Head, a cliffed drumlin, round to a shoal that is all that remains of the former Miseners Head drumlin (Carter et al. 1990). It has

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⊡⊡ Fig. 2.2.14 Gaetz Head, a cliffed drumlin. (Courtesy Geostudies.)

retreated landward as the result of frequent overwash by storm waves (forming washover lobes), and has been breached at its eastern and western ends, where trailing spits have formed. The development of these trailing spits on what is now the Story Head drumlin island and landward from the shoal marking the former Misiners Island drumlin, illustrates the transition from the original swashaligned barrier beach to a series of drift-aligned ridges transverse to the general outline of the coast. Miseners Long Beach is a relatively stable high stormpiled barrier beach of poorly sorted subangular grey pebbles and cobbles behind a sandy foreshore. The barrier beach is steep-fronted and reflective, and often has large beach cusps. Winter shore ice is sufficient to protect the upper beach face and displace the breaker zone seaward. The barrier beach stands in front of Misemers Lake. East of Petpeswick Inlet is Martinique Beach – another long barrier spit with slightly cliffed grassy dunes and a surfing beach. The indented rocky coastline continues north-east to Cape Canso, which shelters Chedabucto Bay. The narrow Strait of Canso separates Cape Breton Island and runs through to St. Georges Bay in the southern Gulf of St. Lawrence. It is crossed by the Canso Causeway, completed in 1955. The south-east coast of Cape Breton Island is again rocky and exposed to Atlantic swell and storm waves. St Peters Island, near Rockdale, has a cuspate spit on its northern (lee) coast, matched by a paired cus­ pate ­spit  at Potheir Point on the mainland. To the east a double tombolo links Michaux Point to the mainland near Gracieville.

The northern coast of Cape Breton Island is strongly influenced by the SW–NE trend of ridges and valleys following Appalachian folds and faults, with intricate bays and straits along submerged valleys. Crystalline rocks form uplands, and lowlands follow corridors of Carboniferous rock, as at Sydney Harbour. Scatari Island is hilly, with minor cliffing on its coasts. Glace Bay has a mean spring tide range of 1.0 m, and Table Head to the north has receding cliffs and isolated stacks bordering a low plateau. Johnson (1925: Fig. 152) provided a set of six photographs showing the dissection of a head­land near Table Head between 1900 and 1921. At Briton Cove there is a good example of a sloping cliff on steeplydipping Carboniferous sandstone, the cliff face coinciding with the bedding plane on some sectors (Johnson 1925: Fig. 83). West of Sydney Harbour Cranberry Head, Merrit Point and Point Aconi have receding cliffs in Carboniferous sediment, including coal seams. At Ingonish a barrier spit extends across South Bay, breached near its southern end, and a long curved barrier spit backs Aspy Bay at Dingwall, which has a mean spring tide range of 0.7 m. Beyond Aspy Bay a small peninsula runs out to Cape North beside Cabot Strait, at the entrance to the >  Gulf of St Lawrence.

4.  Sable Island Lying east of Nova Scotia is Sable Island (44°S, 60°W), an unusual arcuate deposit of unconsolidated sand, 200 km

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south-east of mainland Nova Scotia on the outer margin of the continental shelf. It is only 35 km long, with a maximum width of about 2 km. It has been shaped largely by the prevailing westerly winds and waves in the Gulf of St. Lawrence. Vegetated dunes are found in the central section. Cameron (1965) documented changes in the morphology and size of the end points of the island.

References Amos CL, Long BFN (1980) The sedimentary character of the Minas Basin, Bay of Fundy. In: McCann SB (ed) The Coastline of Canada, Geological Survey Paper Canada 80–10:153–180 Cameron HL (1965) The shifting sands of Sable Island. Geogr Rev 44: 363–376 Carter RWG, Orford JD, Forbes DL, Taylor RB (1990) Morphosedimen­ tary development of drumlin-flank barriers with rapidly rising sea level, Story Head, Nova Scotia. Sediment Geol 69:117–138 Carter RWG, Orford JD, Jennings SC, Shaw J, Smith JP (1992) Recent evolution of a paraglacial estuary under conditions of rapid sea level rise: Chezzetcook Inlet, Nova Scotia. Proc Geol Assoc 103:167–185

Dalrymple RW, Knight RJ, Middleton GV (1975) Intertidal sand bars in Cobequid Bay (Bay of Fundy). In: Cronin LE (ed) Estuarine Research, Vol 2, Academic Press, New York, pp 293–307 Dalrymple RW, Knight RJ, Lambiase JT (1978) Bedforms and their hydraulic stability in a tidal environment, Bay of Fundy, Canada. Nature 275:100–104 Johnson DW (1925) The New England-Acadian Shoreline, Wiley, New York Knight RJ (1980) Linear sand bar development and tidal current flow in Cobequid Bay Bay of Fundy Nova Scotia. In: McCann SB (ed) The Coastline of Canada, Geological Survey Paper Canada 80–10: 123–152 Manson GK (2002) Subannual erosion and retreat of cohesive till bluffs, McNab’s Island, Nova Scotia. J Coast Res 18:421–432 Owens EH, Bowen AJ (1977) Coastal environments of the Maritime Provinces. Maritime Sediments 13:1–13 Shaw J, Ceman J (1999) Salt marsh aggradation in response to late Holocene sea level rise at Amherst Point, Nova Scotia. The Holocene 9:439–451 Taylor RB, Shaw J, Forbes DL, Frobel D (1996) Field Trip Guidebook. Eastern shore of Nova Scotia. Coastal response to sea-level rise and human interference. Natural Resources, Canada

2.3  Gulf of St. Lawrence

Eric Bird

1.  Introduction

coastline is generally fringed by shore ice in winter, and exposed to wave action for up to 8 months in summer: there are variations in the length of winter and summer conditions between the southern and the northern parts. Atlantic ocean swell reaches the south coast of Newfoundland, and to a lesser extent the northeast-facing coasts of the island of Newfoundland and Labrador. Generally the wave regime on these coasts, and in the relatively protected Gulf of St. Lawrence, is dominated by local storm-generated waves.

The coast of the Gulf of St. Lawrence includes the St. Lawrence estuary and the south coast of Quebec (> Fig.  2.3.1). The southern coast of the Gulf is readily accessible and has been studied in detail, but much of the northern coast is rocky, barren and subarctic, difficult of access and little known geomorphologically. The Gulf of St. Lawrence is a microtidal sea environment with a maximum SW-NE extent of 800 km. The

⊡⊡ Fig. 2.3.1 Coastal regions of eastern Canada. 1-Bay of Fundy, 2-Gulf of St. Lawrence (south), 3-Gulf of St. Lawrence (north), 4-St. Lawrence Estuary, 5-Atlantic Nova Scotia and southern Newfoundland, 6-E. Newfoundland (including Labrador). Provinces: A-Quebec, B-New Brunswick, C-Nova Scotia, D-Prince Edward Island, E-Newfoundland. (Courtesy S.B. McCann.)

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2.  Magdalen Islands

The north coast of Prince Edward Island has cliffs cut in glacial drift over soft red Permian sandstone (> Fig. 2.3.3) and beaches that pass west from Stanhope Beach into a chain of barriers, spits and barrier islands with grassy dunes (> Fig.  2.3.4) (McCann 1979). The barriers have developed across embayments and estuaries to enclose lagoons such as Tracadie Harbour and Malpeque Bay, and are basically transgressive, though downdrift, some have prograded dune ridge sequences. Most of the barrier coastline has spring tides smaller than 1.6 m, and tidal currents are only important near inlets. There are short period northeasterly waves generated by occasional storms within the Gulf, which drift sediment westward along the beaches; but the prevailing westerly winds blow from land to sea, and have little effect on the barriers except on sectors with westerly exposure. Tidal inlets that breach the barriers vary in size from small, temporary inlets, only tens of metres wide and 2–3 m deep, to large permanent inlets up to 1 km wide and 15 m deep. They separate barrier islands such as Hog Island which has low grassy dunes, a wide beach with small barchans that move northward to the sea, lumber on the backshore, and Spartina marshes on the lagoon shore, locally eroded to expose peat. Further west a high coast has maintained its dimensions during slow barrier retreat. Longshore transport rates are high at this location, and

Out to the northwest are the Magdalen Islands (> Fig. 2.3.2), where two long barrier beaches and dunes form tombolos joining a series of rock islands and enclosing shallow central lagoons (Owens and McCann 1980). Together with complex terminal spits they make a continuous narrow land area 70 km long. On the west coast of these islands are three sand bars, the outermost 700–1,000 m offshore in a water depth of 4–6 m.

3.  South Coast, Gulf of St. Lawrence West of Cape North, the coast of Cape Breton Island turns southward and faces the Gulf of St. Lawrence. Geologically Cape Breton Island and the north coast of Nova Scotia are part of the northern section of the Appalachian system, and consist predominantly of deformed Palaeozoic sedimentary rocks trending SW-NE. St. Georges Bay opens northward into the Gulf of St. Lawrence and west of Cape St. George, a generally low-lying north coast of Nova Scotia, passes along the southern side of Northumberland Strait, to the south of Prince Edward Island. There are several valley-mouth rias on this microtidal coast.

⊡⊡ Fig. 2.3.2 The Magdalen Islands. (Courtesy Geostudies.)

Bed rock Spits and Barriers Longshore transport

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Gulf of St. Lawrence

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⊡⊡ Fig. 2.3.3 Cliffs cut in Permian sandstone, Orby Head, Prince Edward Island. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.4 Low dunes behind sandy beach, with small tidal inlet at Dalvay Beach, north coast of Prince Edward Island. (Courtesy Geostudies.)

nearshore erosion of the sea floor appears to be contributing new sediment (Armon and McCann 1977). The south coast of Prince Edward Island, facing Northumberland Strait, is embayed, with broad promontories showing cliffs cut in soft Lower Permian clay and sandstone, as at Borden (> Fig. 2.3.5). Charlottetown has a mean spring tide range of 1.8 m, and muddy shores are exposed at low tide. At Gallas Point the receding cliffs are

cut in glacial drift, fronted by broad shore platforms on Lower Permian sandstone (> Fig.  2.3.6) which display weathering and abrasion features, and have a partial cover of marine algae. In the bays and estuaries are salt marshes dominated by Spartina and red muddy intertidal areas. The southern coast of Northumberland Strait in Nova Scotia and New Brunswick shows similar features bordering a broad low plateau of glacial drift. At Cape Jourimain

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⊡⊡ Fig. 2.3.5 Cliffs at Borden on the south coast of Prince Edward Island. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.6 Shore platform at Gallas Point, south coast of Prince Edward Island. (Courtesy Geostudies.)

there are slumping bluffs of till fronted by shore outcrops of almost horizontal Lower Permian pink sandstone exhumed from the drift cover. Sectors of low cliff are cut into the sandstone, as at Cape Bald (Cap Pelée). There is ­extensive nearshore bar topography, as in Kouchibouguac Bay, where there are two bars, the outermost 200–300 m offshore in about water about 2 m deep at low tide

(Davidson-Arnott and Greenwood 1976). North from Point Sapin, beyond the shelter of Prince Edward Island, the coast becomes exposed to stronger wave action across the Gulf of St. Lawrence, and there are spits and barrier beaches (Johnson 1925: 147). Buctouche Bar is a compound recurved spit that has: grown intermittently southward. The mouth of Miramichi Bay is encumbered with

Gulf of St. Lawrence

these depositional forms, with intervening tidal inlets maintained by strong currents, as at Portage Gully (Rein­ son 1977). Portage Island is a barrier island with recurved beach ridges indicating its southward growth. The inlets have large ebb and flood tidal deltas, and have maintained their positions throughout the period of historical record (about 200 years), though there have been changes in the pattern of channels and shoals. Landward transfer of sediment through these low, narrow transgressive barriers of New Brunswick has been by overwash, producing fans of sediment on the landward side, and through tidal inlets, past and present. The beaches and barriers come to an end at Miscou Point, on the southern side of the mouth of the broad Baie des Chaleurs which narrows westward to an estuary at Campbellton. On the northern side is a series of bays and headlands, extending to Bonaventure Island. Beyond this Le Rocher Percé (the Pierced Rock) is an elongated island of steeply dipping to vertical limestone bordered by cliffs up to 100 m high and penetrated by a tall natural arch. A stack at the eastern end is the remains of an earlier archway that collapsed in the nineteenth century. There is a broad gravelly intertidal zone. The Murailles to the west, on the southern side of Mal Bay, are vertical cliffs truncating a ridge of steeply dipping limestone that runs along the coast. They have cut back into the landward slope, and so are receding cliffs of diminishing altitude (Johnson 1925: 200–202).

⊡⊡ Fig. 2.3.7 Strandlines on salt marsh, St. Lawrence estuary. (Courtesy Geostudies.)

2.3

Mal Bay is backed by a barrier spit, and to the north is the Baie de Gaspé, the estuary of the Dartmouth River. Beyond this is the long narrow promontory that ends in Cape Gaspé. Here spectacular cliffs are fronted by shore platforms cut in shale, the outer edge of which has been trimm­ ­ed back by shore ice erosion, particularly when the shore ice is disrupted by spring storm waves (Trenhaile 1978).

4.  St. Lawrence Estuary Cape Gaspé marks the beginning of the St. Lawrence estuary, which has a long curving southern coast backed by the rising slopes of the Shickshock Mountains. Mean spring tide range increases from 1.0 m at Gaspé to 3.7 m at Rimouski and 4.9 m at Quebec as the estuary narrows to the outflow from the Great Lakes (> Great Lakes (Canada)). The intertidal zone widens correspondingly. Shore ice develops and persists for four months in winter, when there is a thick cover of mobile pack ice offshore. On the north coast of the Gaspé Peninsula are seaward sloping intertidal shore platforms cut across dipping strata that run roughly parallel to the coastline. The movement of shore ice to and fro has contributed to the abrasion of these platforms. There are salt marshes (schorre) threaded by high tide strandline litter (> Fig. 2.3.7) and strewn with pebbles, ­cobbles and large ice-­transported boulders with scoured furrows (> Fig. 2.3.8) (Dionne 1972, 1989).

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Boulder barricades are elongated rows of boulders emplaced just offshore on arctic or subarctic coasts by the grounding of boulder-laden ice rafts in the nearshore zone when the ice breaks up in spring. They have been described from the shores of the Gulf of St. Lawrence, notably at Cap la Baleine (Dionne 2003). There are also boulder pavements, assemblages of rocks in the intertidal zone on cold climate coasts flattened by impaction

beneath shore ice, usually smoothed and often striated (Hansom 2005). On the shores of the St. Lawrence estuary erratic boulders have been derived partly from the Appalachian Mountains and partly from the Canadian Shield to the north, and delivered by glaciers. Many boulders have been transported across the estuary from the north shore by drifting ice (Dionne 2002). Shore ice breakup in spring is

⊡⊡ Fig. 2.3.8 Boulder pushed by shore ice, Trois Pistoles, St. Lawrence estuary. (Courtesy Geostudies.)

⊡⊡ Fig. 2.3.9 Displaced raft of salt marsh turf, St. Lawrence estuary. (Courtesy Geostudies.)

Gulf of St. Lawrence

accompanied by erosion as storm waves push it across the salt marsh, displacing rafts of Spartina turf (> Fig. 2.3.9) Abrasion by shore ice has also contributed to the shaping of shore platforms, as at Trois Pistoles (Dionne and Brodeur 1988). At low tide a wide zone of muddy rocks and gravel is exposed in front of the salt marsh. The north coast of the St. Lawrence estuary is more embayed, with outcrops of hard Pre-Cambrian rock along the southern edge of the Canadian Shield and a coastal plain of glacial drift and glacifluvial deltas east from Cap Colombier. Beach sediment are often scarce, and largely derived from the erosion of glacial drift deposits. Emerged coastline deposits are found, particularly along the St. Lawrence estuary. Tide ranges diminish eastward to 2.6 m at Sept Iles, and as the estuary widens wave action becomes stronger. The eastern boundary of the St. Lawrence estuary is arbitrary, but can be taken at longitude 64° 35ʹ where the channels diverge to the Jacques Cartier Passage and the Honguedo Passage, north and south of Anticosti Island.

5.  Anticosti Island Anticosti Island is generally low and hilly, with broad plains of glacial drift and generally smooth coastlines with low cliffs, sand and gravel beaches and salt marshes. A series of 22 terraces planed across local geological structures at up to more than 400 feet above sea level was described by Twenhofel and Conine (1921), who considered them to be of marine origin, a view questioned by Johnson (1925: 144–146). There are separate discussions in the chapters on > Newfoundland and > St. Pierre et Miquelon.

6.  North Coast, Gulf of St. Lawrence The south-facing Quebec coast in the Jacques Cartier Passage and to the east is generally rocky and indented. The shore ice season increases from 3 to 4 months in the St. Lawrence estuary to 6 months in Labrador. An ice foot forms along these coasts in winter, and persists to the breakup season. The offshore ice cover in the south is usually dense, slowly moving pack ice, but in the north there is commonly a broad zone of fast ice extending seaward for many kilometres between the shore ice and the mobile pack. The presence of shore ice for many months each year reduces the period when wave processes are effective, but introduces other effects on the shore, ­particularly at breakup when erosion, transport and deposition by drift ice are key processes. Storm waves occur frequently.

2.3

The Natashquan River flows down to an estuary in a broad deltaic protrusion at Old Post Point, and to the east the low coast has numerous rias between rocky headlands. East of Harrington Harbour it is fringed by numerous small rocky islands. Boulder barricades are found along the mean low tide line and nearshore gravels have been derived from ice-rafted sediment. Wave energy diminishes as the coast passes alongside the Strait of Belle Isle. At Forteau the eastern border of Quebec crosses the coast, and the east coast of Labrador is within the province of > Newfoundland.

7.  Eastern Labrador At York Point on the northern side of the Strait of Belle Isle the Labrador coastline turns northward to Domino, then northwestward. It is a generally steep, rocky coastline, cut by numerous fjords and facing the northwest Atlantic Ocean. The tide range is small (1.9 m at Nain). Storm waves are frequently generated during the passage of cyclonic storms from west to east across the northern Atlantic and the Labrador Sea. There are pocket beaches of sand and gravel, derived from glacial drift, in some of the coves. Boulders on the shore are displaced by ice rafting (Hansom 2005). The effects of glaciation are evident throughout the Labrador region, which was covered by ice as recently as the late Wisconsin. Glacial erosion produced the fjords of eastern Labrador, glacial deposits cover lowlands and much of the shallow continental shelf, and changes in relative sea level due to deglaciation and isostatic rebound are important everywhere. The rocky coast with numerous fjords continues northward to Cape Chidley.

References Armon JW, McCann SB (1977) Longshore sediment transport and a sediment budget for the Malpeque barrier system, Southern Gulf of St. Lawrence. Can J Earth Sci 14:2429–2439 Davidson-Arnott RGD, Greenwood BB (1976) Facies relationships on a barred coast, Kouchibouguac Bay New Brunswick. In: Davies RA, Ethington RL (eds) Beach and nearshore sedimentation. Society of Economic Paleontologists and Mineralogists, Special Publication 24:149–168 Dionne JC (1972) Caracteristiques des schorres des regions froides, en particulier de 1¢estuaire du Saint Laurent. Z Geomorphol 15:137–180 Dionne JC (1989) The role of ice and frost in salt marsh development. A review with particular reference to Quebec, Canada. Earth Sci Rev 26 Dionne JC (2002) Dolomite erratics at Rivière-Blanche, south coast of the maritime estuary of the St Lawrence (in French). Can J Earth Sci 39:1239–1255

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Dionne JC (2003) The boulder barricade at Cap la Baleine, north shore of Gaspé; peninsula (Québec): nature of boulders, origin and significance. J Coast Res 18:652–661 Dionne JC, Brodeur D (1988) Erosion of rocky platforms by shore ice (in French). Z Geomorphol 32:101–115 Hansom JD (2005) Boulder pavements. In: Schwartz ML (ed) Encyclopedia of Coastal Science, Springer, Dordrecht, pp 208–210 Johnson DW (1925) The New England-Acadian shoreline. Wiley, New York McCann SB (1979) Barrier islands in the southern Gulf of St. Lawrence. In: Leatherman SP (ed) Barrier Islands from the Gulf of St Lawrence

to the Gulf of Mexico. New York, Academic Press, pp 29–63 Owens EH, McCann SB (1980) The coastal geomorphology of the Magdalen Islands, Quebec. In: McCann SBG (ed) The Coastline of Canada, Geol Surv Can Pap 80–10:123–152 Reinson GE (1977) Tidal current control of submarine morphology at the mouth of the Maramichi estuary, New Brunswick. Can J Earth Sci 14:2524–2532 Trenhaile AS (1978) The shore platforms of Gaspé. Quebec. A Assoc Am Geogr 68:95–114 Twenhofel WH, Conine WH (1921) The post-glacial terraces of Anticosti Island. Am J Sci 1:268–278

2.4 Saint-Pierre-et-Miquelon Islands Jean-Marie M. Dubois 

1.  Introduction The Saint-Pierre-et-Miquelon archipelago is a small French territory, only 242 sq. km, located 22 km from the southern coast of Newfoundland, between 46° 45' and 47° 10' N and 56° 05' and 56° 25' W. It constitutes an enclave on the Saint-Pierre Bank within Canadian territorial waters. The archipelago is formed by a main island (Saint-Pierre) to the east, where the prefecture and administrative centre are located, by two more important islands (Miquelon and Langlade, joined by a N-S tombolo) as well as six other islands (Le Cap, Grand Colombier, Verte, aux Marins, aux Vainqueurs and aux Pigeons islands) and various islets and reefs (Aubert de la Rue 1970) (>Fig. 2.4.1). The Labrador Current, which flows through the archipelago from east to west, is responsible for cool and foggy summers (85–120 days) and the average annual tem­ perature in this subarctic marine environment is barely 5.5°C. Average annual precipitation is about 1,424 mm – about a third as snow. Winds exceed 60 km/h for almost half the year, and storms are frequent. The storm of January 1987 caused up to 5 m of recession on cliffs cut in glacial drift deposits, the February 1988 storm, with winds of up to 160 km/h cut back cliffs by up to 14 m and the storm of December 1989 caused 7 m of erosion (Rabottin 1990). Much of the archipelago has scrubby vegetation following long-term deforestation, but woodland occupies 20% of Langlade Island. Peatland is extensive, and heathland and grassland are found over coastal sand and pebble areas. The main algae found along the coasts are laminaria, brown algae, red algae and lime-secreting algae. The geology is varied (Dubois 2006). Le Cap Island is  mainly composed of metamorphic crystalline PreCambrian rocks (schists, quartzites, paragneiss and migmatites), with a small intrusive granitic stock. Miquelon Island consists of slightly metamorphosed post-­Ordovician volcanic rocks, mainly rhyolites with some basalts, andesites and breccias. Langlade Island has slightly meta­morphosed Appalachian Palaeozoic sedimentary rocks (sandstones, schists, phyllites, limestones, quartzites and conglomerates) trending SW-NE with faults. Saint-Pierre and adjacent islands are composed of post-Ordovician intrusive rocks, mainly rhyolites, brecchia and tuffs with a few andesites.

The islands are mostly rocky with superficial deposits of glacial till, wind-blown sand, alluvium, weathered rock, landslide debris, colluvium and peat. The islands are the remains of an ancient erosion surface except for Langlade Island, which has a tabular relief. On the main islands there are a number of summits: Morne de la Grande Montagne on Miquelon island (240 m), Le Trépied on Saint-Pierre island (207 m), Tête de Cuquemel on Langlade Island (190 m) and a hill with an altitude of 153 m at the NE end of Le Cap island. The thick superficial deposits are located mainly on Miquelon and Langlade Islands, while silty clay till, superficially reworked by the sea, is present over most of the periphery of Miquelon island and in NW Langlade Island, often affected by landslides. Tombolos of sand and gravel link Langlade and Miquelon Islands, Miquelon and Le Cap Island, and will eventually link aux Vainqueurs and aux Pigeons islands. The archipelago was covered by ice from Newfoundland during the Middle Wisconsin, as shown by the nature of the erratics, but not during the Late Wisconsin. Glaciation left moraines to the south of Saint-Pierre Island and till over most of the islands. A subsequent marine submergence, around 38,000 years bp left shore deposits and indications of marine erosion up to at least 25 m above sea level. The southern limit of the last glaciation was along the south coast of Newfoundland (Piper and Macdonald 2001), and evidence of periglaciation – including frost wedges – is found on all the islands (Lauriol and Dubois 1981). However, recent shore formations at 10 m on the north of Miquelon Island, 6.4 m in the north of Langlade Island and present sea level on Saint-Pierre Island, could indicate differential Post­ glacial emergence. At Quine Point, north of Grand Barachois, the base of a peat bog, 1 m below present sea level, was dated at 8,460  ±  150 years bp, indicating recent submergence (Lauriol and Dubois 1981). The base of another peat bog close to present sea level in the De Savoyard cove, on the south coast of Saint-Pierre Island, had an age of 2,000 ± 60 years bp. Global warming is expected to cause submergence of about 2 mm/year, taking account of continuing glacioisostatic subsidence (shaw and Forbes 1990). In addition to wave action there are longshore currents which have contributed to the shaping of the tombolos on Langlade, Miquelon and Le Cap Islands (Dubois 1980).

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⊡⊡ Fig. 2.4.1 The Saint-Pierre-etMiquelon archipelago.

Generally, the longshore currents flow in a clockwise direction, but there are local reversals in the vicinity of large capes on the eastern shore of the Langlade isthmus and the tombolo between Miquelon and Le Cap Islands. Sediment are certainly carried offshore by convergent longshore currents to the west of Miquelon Island. There is historical evidence of coastal erosion in the archipelago since the eighteenth century, but there have

also been local phases of accretion, as on the Langlade isthmus. Using landmarks established between 1984 and 1989 and comparisons of the 1952 and 1985 aerial photographs, Rabottin (1990) concluded that 88% of the coasts in superficial deposits on Langlade and Miquelon Islands had receded at an average annual rate of slightly more than 1 m, and that this rate was increasing. On lagoon barriers, erosion sometimes reached 2–3 m/year. On the morainic

Saint-Pierre-et-Miquelon Islands     

shores of Miquelon Island the annual recession rate is smaller, between 0.1 and 0.4 m. Further­more, according to fishermen, it seems that during the last 80 years winters have been less rigorous and the winter ice-foot offers less littoral protection than before (Lauriol et Dubois 1981). There are only nine areas of accretion, all on the Langlade isthmus and the tombolo between Miquelon and Le Cap Islands. Rabottin (1990) found that the average rate of accretion was slightly more than 1 m/year. Accretion occurs near sediment sources such as eroding cliffs and around lagoon outlets. The rocks of the cliffs are relatively friable and are disintegrated by cycles of freeze-and-thaw, after which waves remove the fallen material.

2.  Island Coastlines Le Cap Island is linked to Miquelon Island by a sand and pebble tombolo more than 3 km long, and in places, less than 100 m wide. This double tombolo, which was built before the sixteenth century, contains a lagoon with an ­outlet on the eastern side. Since its altitude is only 1–4 m, the tombolo is often inundated during storms coinciding with high tides. During the winter of 1947–1948 a 2.5 m rampart of ice and pebbles was pushed up by storm waves, and this was still visible at the end of May. Le Cap Island has an indented coast with high rocky cliffs which are eroding. Miquelon Island has low morainic and slightly in­­ dented coasts with numerous eroding peat bogs at times. On the west coast cliffs cut into till have a height of 5–10 m, and are receding (Lauriol et Dubois 1981), but a few rocky points and reefs protrude. On the lower east coast some cliffs are rocky, but numerous beach barriers of sand and pebbles enclose lagoons of varying size. The Langlade isthmus is a 12 km long and 100 m to 6 km wide tombolo composed of sand and pebbles with an altitude varying from 1–4 m, and dunes west of Grand Barachois reaching an altitude of 26 m. Captain James Cook mapped an opening about 1.8 km wide in the isthmus in 1763, but this has been closed since 1784. There are parabolic dunes 10–20 m high (locally named buttereaux) at both ends, where spits shelter two large lagoons with outlets towards the east. The west coast is strewn with shipwrecks. It is estimated that there have been more than 650 shipwrecks around the archipelago since 1816 and this coast has become known as the ship necropolis (Sanguin 1983). About 20% of these wrecks occurred between the end of August and the beginning of October, close to the Fall Equinox (Dubois 1980). Langlade Island has a high plateau with an altitude of 130–140 m, incised by many valleys with permanent water

2.4

courses, notably the Belle Rivière – the only real river in the archipelago. The island is surrounded by high receding rocky cliffs up to 25 m high, except in the north where the coastline is lower. This low coastline, cut mainly in till and glacio-fluvial deposits, with recent shore deposits to the NW, is mostly receding (Dubois 1980). In the north and west of Saint-Pierre Island, the coasts are rocky and rectilinear, with cliffs up to 75 m high, sometimes receding. In the south the coast is low, rocky and indented, with reefs and numerous coves containing pebble and boulder beaches (Mailhot et Dubois 1983). There are only a few sandy beaches, notably at De Savoyard and Alumette, where unfortunately sand extraction occurred in the past. Fortunately this practice has been prohibited since 1984 (Rabottin 1990). A fault zone 50–100 m wide may separate the two types of coast. Saint-Pierre Island is separated from Langlade Island by a strait often referred to as La Gueule de l’Enfer (the Mouth of Hell) because of its violent currents. Breakwaters have been built to shelter the harbour at Saint Pierre. Since 1974 breakwaters have also been built to counter erosion in the inner parts of the coves, but erosion is widespread on the coasts of the smaller islands (Dubois 1980).

References Aubert de la Rüe, E (1970) Saint-Pierre et Miquelon. Horizons de France, Paris, 174 Dubois JM (ed) (1980) Géomorphologie des îles Saint-Pierre et Miquelon: premier rapport d’étape. Département de Géographie, Université de Sherbrooke, 47 Dubois J-MM (2006) Saint-Pierre-et-Miquelon: un laboratoire de terrain en miniature pour les géographes physiques. Enjeux géographiques 2:36–44 Lauriol B, Dubois JM (1981) Géomorphologie des îles Saint-Pierre et Miquelon, deuxième rapport d’étape: le Grand Barachois et la côte ouest de l’île de Miquelon. Département de géographie, Université de Sherbrooke, 61 Mailhot R, Dubois JM (1983) Géomorphologie des îles Saint-Pierre et Miquelon, troisième rapport d’étape: classification de la zone côtière des îles Saint-Pierre et Langlade. Département de géographie, Université de Sherbrooke, 40 Piper DJW, Macdonald A (2001) Timing and position of Late Wisconsinan ice-margins on the upper slope seaward of Laurentian Channel. Géographie physique et Quaternaire 55:(2):131–140 Rabottin JL (1990) Saint-Pierre et Miquelon, un archipel d’outre-mer assiégé par les flots. Cahiers Nantais 35–36:37–47 Sanguin A-L (1983) Saint-Pierre et Miquelon. Département français d’Amérique du Nord, Norois, Poitiers, p 110 Shaw J, Forbes DL (1990) Short and long-term relative trends in Atlantic Canada, National Research Council of Canada. Proceedings of the Canadian Coastal Conference 1990, Kingston, Ottawa, Canada, pp 291–305

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2.5  Newfoundland

Eric Bird

1.  Introduction The province of Newfoundland and Labrador includes the large island of Newfoundland (405,720 sq. km). This chapter deals with the island of Newfoundland and the Labrador coast up to Cape Chidley. Much of the coastline is rocky, barren and subarctic in character.

2.  Newfoundland The island of Newfoundland partly blocks the eastern end of the Gulf of St. Lawrence and is separated from Labrador to the north by the Strait of Belle Isle and from Cape Breton Island to the southwest by the wider Cabot Strait. It consists of Appalachian structures, with Palaeozoic rock formations, folded and faulted mainly along SW–NE alignments, in contrast with the Laurentian Shield structure of Labrador. The interior consists of a broad plateau with low hills and a higher ridge (Long Range Mountains) runs NNE along the western peninsula. The rugged island coastline is about 16,000 km long, with peninsulas and bays following the SW–NE trend, and much of the coast is backed by hills or low mountains, except for the lowlands in the northeast, where Pleistocene glaciation has left numerous fiords between narrow peninsulas. There are fiords on the steeper western and southern coasts. The Newfoundland coast is rising as the result of postglacial isostatic recovery. Coastline changes have, however, been generally very slow and the outline charted by Captain Cook in 1775 has been little modified. Newfoundland has cold winters and cool to mild summers. Shore ice is present along much of the seaboard in winter, only the outer coasts of southern Newfoundland being ice free, although even there ice may form in sheltered bays. The Strait of Belle Isle has sea ice for up to 10 months each year and is subject to strong tidal currents. Storms and strong winds are frequent, and the coast is often foggy. Westerly winds over the Gulf of St. Lawrence produce waves that generate northward drift along the west coast, where beaches, spits and barriers of sand and gravel have been derived from cliffs cut in glacial drift. Cuspate

spits have formed on the lee of islands, one of which (Cow  Head) is attached by a tombolo. Coastal dunes ­drifting inland have overrun a forest near Western Brook in the Gros Morne National Park. Long Point and St. George’s are major spit formations. Atlantic Ocean waves reach the southern and eastern coasts, which are generally steep and cliffy. There are many small islands off the coast, including the French archipelago, > St. Pierre et Miquelon, off the south coast. To the southeast is the submerged Grand Banks, sustaining major fisheries.

3.  Labrador At York Point on the northern side of the Strait of Belle Isle the Labrador coastline turns northward to Domino, then northwestward. It is a generally steep, rocky coastline, cut by numerous fiords and facing the northwest Atlantic Ocean. Many hundreds of kilometres of coastline are difficult to reach and remain little explored in terms of coastal geomorphology. The tide range is small (1.9 m at Nain). Storm waves are frequently generated during the passage of cyclonic storms from west to east across the northern Atlantic and the Labrador Sea. There are pocket beaches of sand and gravel, derived from glacial drift, in some of the coves. Boulders on the shore are displaced by ice rafting (Rosen 1979, 2005). The effects of glaciation are evident throughout the Labrador region, which was covered by ice as recently as the late Wisconsin. Glacial erosion produced the fiords of eastern Labrador, glacial deposits cover lowlands and much of the shallow continental shelf, and changes in relative sea level due to deglaciation and isostatic rebound are important everywhere. The rocky coast with numerous fiords continues northward to Cape Chidley.

References Rosen P (1979) Boulder barricades in central Labrador. J Sediment Petrol 49:1113–1123 Rosen P (2005) Boulder barricades. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 204–206

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

2.6  Northern Canada

Wayne Pollard

1.  Introduction The northern coast of Canada is roughly 172,950 km long, within the Arctic (Shaw et al. 1998). The Arctic is marked by extremely cold winters and year-round freezing air temperatures that are together responsible for widespread permafrost and glaciers as well as the seasonal and perennial occurrence of sea ice and snow. The Arctic is also the traditional home for roughly 45,000 Inuit, the only indigenous population residing in 35 coastal communities. The aim of this discussion is to describe the nature and pattern of the northern coastline of Canada from Battle Harbour in Labrador (Labrador Sea) to Demarcation Point at the Yukon–Alaska border (Beaufort Sea), including the Arctic Archipelago, Ungava Bay, Hudson Bay and the northern mainland. A modified version of Owens’ (1994) regional coastal framework is used to subdivide this region into 11 coastal zones (>Fig. 2.6.1). These zones reflect physiographic patterns linked mainly to bedrock geology and location, from crystalline Pre-Cambrian (Archean and Proterozoic) rocks in Labrador, Ungava Bay, eastern Baffin and Ellesmere Islands, Palaeozoic sedimentary rocks in the central Arctic Archipelago and southern Hudson Bay to poorly consolidated Mesozoic and Cenozoic sediment along the north-westerly edge of the Arctic islands and the Beaufort Sea (Bird 1967). The geomorphic diversity of these northern coastal environments includes deltas, estuaries, lagoons, barrier islands, spits, fringing sand and gravel beaches, steep bedrock cliffs, and several systems unique to northern environments such as fiords, ice coasts (tidewater glaciers and ice shelves), ice-rich permafrost coasts, breached thawlake coasts, and thermokarst coasts. The complexity of these systems is increased by regional patterns of isostatic adjustment and wave climate. Sea level rise and littoral processes related to changing storm patterns and changing sea-ice conditions add to the diversity. The dynamic nature of any coastal system is a function of the physical and morphologic characteristics of the coastline materials as well as the environmental forces operating on it. There are dramatic contrasts in tide range. According to Shaw et al. (1998), the tide range decreases northward

and westward across the Arctic Archipelago. The largest range occurs in Frobisher Bay and Ungava Bay where it is about 13 m. A tide range of more than 4 m occurs in both southern (Hudson Strait) and northern Baffin Island (Nares Strait), while tides are less than 1 m in east central Baffin Island. Tides are more than 4 m in the Foxe Basin, 1–2 m in Barrow Strait and typically less than 0.5 m along the western and northern edge of the Arctic Archipelago and the Beaufort Sea. In Hudson Bay, the tide range is less than 4 m. In some areas, for example, the Beaufort Sea, storm surges can add 1–2 m of water-level rise to the tides. Relative sea level change is determined from either tide gauge records or calculated from coastal emergence curves. In the Canadian Arctic, there are only two locations where tidal records are long enough to detect current trends. At Tuktoyaktuk, tide records indicate a relative rate of sea level rise (submergence) of about 4 mm/year for the Beaufort Sea area. By comparison, tide gauge records from Churchill indicate an emergence rate of the order of 8–9 mm/year for western Hudson Bay. Emergence curves for the rest of the Canadian Arctic suggest emergence rates of up to 7 mm/year across most of the central Arctic including Hudson Bay and submergence of the order of 5 mm/year for eastern Baffin Island. The geomorphic evidence for emergence includes raised beach and delta deposits, marine sediment, marine shells, fossils and driftwood. Under current scenarios of climate change, many Arctic coastlines and several communities will be susceptible to flooding.

2.  The Coastlines Canada’s northern coasts are divided into regions based on common physiographic character and geology. The Labrador Coast (Zone 1) is divided into southern and northern segments. The southern segment from Battle Harbour to Nain consists of roughly 600 km of rocky coastline dominated by numerous small bays and inlets, rugged headlands and islands. The hard crystalline nature of the Pre-Cambrian rocks south of Lake Melville maintain low bluffs and cliffs and a steep nearshore

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 2.6.1 Map of northern Canada showing coastal regions based on common geology and physiography.

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COASTAL SUBDIVISIONS Labrador Hudson Strait-Eastern Baffin Is. Hudson Bay and James Bay East Baffin Bay to Lincoln Sea Jones, Lancaster and Prince Regent Sounds West Ellesmere and Axel Heiberg Islands Northern Ellesmere Island Northwest Arctic Archipelago Central Arctic Archipelago Southern Arctic Archipelago and Mainland coast Beaufort Sea

bathymetry that limit the development of littoral deposits and landforms to a few sheltered bays. The Groswater– Lake Melville area contains unconsolidated materials that form extensive beaches, beach ridges and coastal dunes. Small estuaries occur in some of the sheltered bays. Cape Harrison to Nain displays a mix of rugged and steeply sloping rocky coasts with small bays and headlands. In the Nain area, low relief and gentle nearshore slopes develop wide boulder-strewn tidal flats (>Fig. 2.6.2), bolder barricades and shingle beaches.

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The north segment from Nain to Cape Chidley is roughly 450 km of rugged coastline dominated by the Torngat Mountains that rise abruptly to elevations over several hundred metres. The hard igneous rocks of the Canadian Shield (Nain Province) have been glacially eroded producing a coastline dominated by numerous large bays, islands and several deep fiords. For example, Nackvak and Saglek Fiords extend 20–40 km into the Torngat Mountains. In general, this section of the coast is marked by a paucity of beaches and tidal flats except for

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⊡⊡ Fig. 2.6.2 Boulder-strewn tidal flat along the Labrador coast. (Courtesy P. McLaren, Geological Survey of Canada.)

several small deltas that form at the head of fiords and in larger bays. Hudson Strait–Southeastern Baffin Island (Zone 2) is characterised by moderate-to-high relief coastlines underlain by hard Proterozoic crystalline rocks. Ungava and Frobisher Bays are known for their large tidal range and wide intertidal mudflats with boulders. The south side of Hudson Strait has a steep cliff coastline with a few bays and fiords. The north side is irregular, rugged and rocky with several large bays (e.g. Frobisher Bay) and fiords. The coast of Ungava Bay is low and rocky on both the east and west sides, while the south coast is low with wide intertidal mudflats and estuaries (e.g. George River). Small bays on the west side of Ungava Bay also display wide tidal flats. The Hudson Bay–James Bay area (Zone 3) fronts a large (1.23 million square kilometres), shallow inland sea that is covered with ice for 8–10 months of the year. The east coast of Hudson Bay is characterised by gently dipping (seaward) sedimentary strata forming a series of low cuesta islands parallel to the shore and farther offshore (e.g. Belcher Islands). The coast is gently sloping with low beaches and narrow tidal flats. The east coast of James Bay is characterised by Archean crystalline rocks mantled by till, glacial fluvial and marine (Tyrell Sea) deposits. The irregular coastline consists of numerous low rock outcrops, sand and gravel beaches, tidal flats and salt marshes. The coast on the west side of James and Hudson Bays is dominated by several estuaries with wide salt marshes, separated by coastal segments with tidal flats with sand

and gravel ridges oriented normal to the shore and ­shore-parallel sand-gravel beach ridges fronted by tidal flats and salt marshes. The extensive estuarine coastal systems reflect the importance of the Hudson Bay as the destination for many drainage systems in central mainland Canada. The northern part of the west side of Hudson Bay, including the south side of Southampton and Coates Islands, is characterised by low sandy beaches and wide tidal flats. The East Baffin Bay to Lincoln Sea (Zone 4) includes the western coast of Baffin, Devon and Ellesmere Islands. The coasts in this zone are underlain mainly by hard Pre-Cambrian crystalline rocks except for the Eclipse Sound area of northeastern Baffin Island where there are Cretaceous and Tertiary shales and sandstones, and parts of eastern Devon and northern Ellesmere Islands where there are pockets of Cambrian–Ordovician sedimentary rocks (Shaw et al. 1998). Unconsolidated Quaternary sediment are also locally important. All three islands are mountainous and heavily glacierized with active glacial systems that impinge on this coast, including ice caps, outlet glaciers and tidewater glaciers. The coastal relief is generally high and incised by long, winding steep-sided and talus-banked fiords (>Fig. 2.6.3) up to 98 km long. This zone has Canada’s greatest concentration of fiords (Syvitski et al. 1987). Unconsolidated shores are restricted to low coastal forelands and fiord heads where there are extensive sandur and deltaic deposits. In these areas, coastal landforms include low coastal bluffs, broad tidal flats with boulder barricades, fringing and barrier beaches, and

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⊡⊡ Fig. 2.6.3 Long and twisting fiord from northeastern Baffin Island. (Courtesy R. Taylor, Geological Survey of Canada.)

raised beaches and deltas. The Cape Ashton area has an extensive beach–barrier island complex with aeolian dunes (Taylor et al. 1987). Bylot Island displays a range of barrier beach and nearshore bar complexes. Ice coasts comprise less than 1% of the coastlines and are most common along parts of Devon and southeastern Ellesmere Islands. The Jones, Lancaster and Prince Regent Sound (Zone 5) coastal zone includes northwestern Baffin Island, Devon Island, southern Ellesmere Island, southeastern Cornwallis Island and eastern Somerset Island. This region is underlain by relatively weak, flat-lying Palaeozoic limestone and dolomite. It is described by Shaw et al. (1998) as ‘a plateau region, deeply dissected by fiords. Scree-banked coastal cliffs 100–600 m high dominate the coastline except where the coastal plain is sufficiently wide to allow the formation of depositional beach features.’ Sea-ice conditions are considered severe and result in ice scour and ice push in the littoral zone. During the relatively short open water period (Fig. 2.6.4). Locally, bay and fiord head deposits include typical delta, low beach and intertidal flats. The West Ellesmere and Axel Heiberg Islands (Zone 6) zone lies on the eastern edge of the Sverdrup Basin, and is characterised by poorly consolidated folded Mesozoic and Cenozoic sedimentary rocks intruded by more resistant volcanics. Eureka Sound, the narrow passage that separates the two islands, is bordered by extensive lowlands composed of poorly consolidated Cenozoic sediment mantled by a veneer of marine sediment. Several long fiords (Canon, Tanquary, Gibbs, Mokka, Slidre and Greely Fiords) and bays feed into Eureka Sound. The coastline is made up of several straight segments characterised by: (a) gentle slopes with narrow sand and gravel beaches, (b) low-to-moderate bedrock bluffs fronted by narrow beaches, and (c) bays and fiords characterised by welldeveloped sand and gravel beaches, deltas, glacial flood plains, and in places, raised beaches and deltas. A few tidewater glaciers provide a continuous supply of small icebergs that become stranded in shallow bays. The south and west sides of Axel Heiberg Island are dominated by several large steep-sided fiords (>Fig. 2.6.5) (Li, Middle, Strand, Expedition, Glacier, Wolf and Skaare Fiords) and bays (Good Friday, Sand and Whitsunday Bays). Run-off from two large ice caps (Müller and Stacie Ice Caps), several smaller highland ice caps and valley glaciers (ice covers roughly 31.5% of the island) contribute to extremely high sediment loads that form large sand and gravel flood plains at the head of most fiords. Fiord sides are generally

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⊡⊡ Fig. 2.6.4 Gravel beach backed by raised beach ridges, Somerset Island.

⊡⊡ Fig. 2.6.5 Fiord head with a fluvial delta and glacial flood plain, west Axel Heiberg Island.

steep with narrow beaches, or vertical with no beach. Low coastal plains with sand and pebble beaches, raised beaches and muddy lagoons occur in areas separating fiords on the west side of the island. The presence of sea ice year round limits littoral processes but contributes to localised ice push.

Northern Ellesmere Island (Zone 7) includes some of the most northerly coasts in the northern hemisphere and is considered an ice coast. Northern Ellesmere Island is underlain by Palaeozoic and Proterozoic rock covered by a veneer of Quaternary sediment and ice. The Grant Land Mountains is a high, ice-capped fold sequence responsible

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for the high relief (up to 1,200 m) coastal landscape dissected by ice-choked fiords, epishelf lakes and ice shelves. The absence of open water results in a paucity of littoral processes and landforms. Over the past century, the once near-continuous ice shelf that rimmed most of northern Ellesmere Island has disintegrated into a series of smaller ice shelves (Jeffries 1987), which in recent years have almost disappeared (e.g. Ward Hunt and Ayles ice shelves). Northwest Canadian Arctic Archipelago Coast (Zone 8) consists of a series of low-relief (Fig. 2.6.6). They are mostly low-lying, generally Fig. 2.6.7) are also common, especially along the Yukon Coastal Plain and the Tuktoyak­tuk Peninsula. Wetlands, tidal flats and delta deposits characterise the mouth of the Mackenzie River and a number of other smaller rivers. Other landforms include submerged tundra (>Fig. 2.6.8) and breached thermokarst lakes, which are widespread along the Tuktoyaktuk

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⊡⊡ Fig. 2.6.9 Erosion of ice-rich unconsolidated sediment, Beaufort Sea coast.

Peninsula. This coast is erosional in nature with an average annual retreat of about 1 m/year and a maximum annual retreat of 22.5 m/year (Lantuit and Pollard 2008). Areas of exposed massive ground ice develop retrogressive thaw slumps and thermokarst exposures (>Fig. 2.6.9). Dallimore et al. (1996) have shown that degradation of sub-sea permafrost in the nearshore also contribute to shore face as well as cliff retreat. This coast has been identified as highly vulnerable to the effects of sea level rise and global warming (Shaw et al. 1998). Storm waves and storm surges play an important role in coastal dynamics.

References Bird B (1967) The physiography of Arctic Canada. John Hopkins Press, Baltimore, MD, 336 pp

Dallimore SR, Wolfe SA, Solomon SM (1996) Influence of ground ice and permafrost on coastal evolution, Richards Island, Beaufort Sea Coast, N.W.T. Can J Earth Sci 33:664–675 Jeffries M (1987) The growth, structure and disintegration of Arctic ice shelves. Polar Record 23:631–649 Lantuit H, Pollard WH (2008) Fifty years of coastal erosion and retrogressive thaw slump activity on Herschel Island, southern Beaufort Sea, Yukon Territory, Canada. Geomorphology 95:84–102 Owens E (1994) Canadian Coastal Environments, Coastline Processes, and Oil Spill Cleanup Environmental Protection Series Report 3/SP/5 Environment Canada, Ottawa, 328 pp Shaw J, Taylor R, Forbes M, Solomon D (1998) Sensitivity of the coasts of Canada to sea level change. Geol Survey Canada Bull 505:79 Syvitski J, Burrel D, Skei J (1987) Fjords: processes and products. Springer Verlag, New York, 379 pp Taylor R, Praeg D, Syvitski J (1987) Coastal morphology and sedimentation, Eastern Baffin and Bylot Islands, N.W.T. In: Sedimentology. Experiment Data Report, Vol 3. Geological Survey of Canada Open File 1589

2.7 Great Lakes (Canada)

Mary-Louise Byrne

1.  Introduction The Great Lakes are five interconnected freshwater lakes in central eastern North America. Closest to the continental interior is Lake Superior, which flows through the St. Mary’s River into Lakes Michigan and Huron. These lakes empty into the connecting channels of the St. Clair River, Lake St. Clair and the Detroit River that in turn flow into the southernmost of the Great Lakes, Lake Erie, which empties into the Niagara River and Welland Canal, and then Lake Ontario. Finally, Lake Ontario empties into the St. Lawrence River. Overall, the Great Lakes span more than 1,200 km from the head of Lake Superior to the St. Lawrence River. According to the Great Lakes Atlas (2000), the lakes contain about 23,000 km3 of water covering 244,000 km2 of area. They have a total coastline length of more than 17,000 km in both Canada and the United States. The drainage area that empties into the lakes covers more than 521,000 km2. The climate of the Great Lakes basin is greatly influenced by its location in the continental interior (Carter and Haras 1985). Weather patterns are complex and change rapidly because of the location of the basin at the boundary of the polar and tropical air masses. The lakes modify temperature and humidity by acting as heat sources in winter and heat sinks in summer. Precip­­ itation  ranges considerably in this area. For example, mean annual precipitation varies from a total of 74.6 cm (55.9 cm rainfall and 18.7 cm snowfall) in Thunder Bay in  the northeast, to 89.3 cm (80.4 cm in rainfall and 8.9 cm in snowfall) in Kingsville in the south, to 100.1 cm (79.4 cm in rainfall and 20.7 cm in snowfall) in Cornwall in the southeast. The prevailing westerly winds create surface waves and currents in the Great Lakes, and can cause short-term fluctuations in lake levels during intense storms. The influence of these variations is greatest in Lake Erie, the shallowest of the lakes. There are also seasonal and longer term fluctuations in lake levels. Long-term fluctuations are influenced by variations in the balance of precipitation and evaporation. Variations between record high levels

and record low levels range from 1 to 2 m. The greatest variations are on the middle lakes (Erie, Huron and Michigan) and lowest on Lake Superior (the largest) and Lake Ontario (the most controlled). The coastlines of the Canadian side of the lakes ex­­­ hibit great variability from Lake Superior through to the St. Lawrence. The Upper Great Lakes consist of Lake Superior, Lake Michigan and Lake Huron. Lakes Superior and Huron have coastlines in Canada, but Lake Michigan lies wholly within the United States and will therefore not be included here. Geologically, the basins of these lakes appear to have their origins in the Pleistocene glaciations of the last 2.5 million years (Eyles 2002). Our knowledge of their geology prior to that time is very limited. The ice sheets that advanced over Canada and the northern United States deeply eroded river valleys producing the basins that the present lakes fill. During glacial retreat, the lakes had several configurations, approximating the current five lakes and connecting channels by about 8,000–10,000 years bp. The exposed shield rock in the northern portions of the Upper Lakes shows evidence of erosion during glaciation, while the deep Pleistocene deposits of the Lower Great Lakes dominate the present coastlines.

2.  Lake Superior Lake Superior, the most northerly and least heavily settled of the Great Lakes, lies at 183 m above sea level and is shaped by the mid-continent rift system. Pre-Cambrian rock of the Canadian Shield is exposed on its northern coastline (Eyles 2002). Total coastline length in Canada is 1,394 km with an additional 990 km on islands. The coast is dominated by rock outcrops with a thin cover of glacial deposits, with glacial lake sediment close to the shore, and unstratified ground and end moraine further inland. The present coastline consists of high cliffs with coarsegrained pocket beaches, punctuated by deep bays like that of Thunder Bay and Nipigon Bay (Carter and Haras 1985).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_2.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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3.  Lake Huron Lake Huron (176 m above sea level) has the longest coastline on the Canadian side, more than 2,044 km on the mainland and 2,768 km on the many islands that dot the lake and its large bays. Lake Huron is separated from Lake Michigan by the wide Straits of Mackinaw, and the two lakes are considered hydrologically one. Water flows from the Straits to the main basin of Lake Huron. To the north is Manitoulin Island, separated from the mainland by the North Channel. Manitoulin Island is the world’s largest freshwater island, and is composed of low-sloping limestone coast on the main lake side and steeper cliffs on the North channel side. The north coast of North Channel and the east coast of Georgian Bay are composed of Pre-Cambrian rock, with rare pocket beaches of sand as in Killbear Provincial Park. The eastern side of the Bay is known as the 30,000 Islands. Further south in Georgian Bay, the eastern coast becomes more gently sloping, sand is more abundant, and features like Wasaga  Beach in Nottawasaga Bay have accumulated. This area marks the transition from the Pre-Cambrian basement to the overlying Ordovician, and sub­ sequently  the Devonian and Silurian Rocks of the Niagara Escarpment that outcrop on the western side of

Georgian Bay. Spectacular limestone and dolostone cliffs plunge into the waters of the Bay, the water being more than 160-m deep immediately off the northern tip of the Bruce Peninsula (>Fig. 2.7.1). The cliffs have been dissected along faults and joints (>Fig. 2.7.2). The Lake Huron side of the Bruce Peninsula is characterised by the same gently sloping limestone seen in Manitoulin Island. In the north, the bedrock is often exposed (>Fig. 2.7.3). Moving south along the Lake Huron coast, steep bluffs of Pleistocene till are common. The easily eroded till supplies a wealth of sand that has accumulated in the southern end of a littoral cell at Grand Bend and Pinery Provincial Park, dominated by wide low-­sloping sand beaches and dunes (Dech et al. 2005) (>Fig. 2.7.4). Ice plays a role in the transport of sediment (>Fig. 2.7.5), and waves and currents generate sediment movement in the ice-free season (Houser and Greenwood 2005).

4.  Lake Erie Moving through the connecting channels of the St. Clair River, Lake St. Clair and the Detroit River, the transition is made from the Upper to the Lower Great Lakes. Lake Erie (173 m above sea level) has a coastline length in Canada

⊡⊡ Fig. 2.7.1 Cliffs and shore platforms with cobble and gravel beaches on the east side of the Bruce Peninsula, Lake Huron. Cobbles are derived from the cliffs.

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⊡⊡ Fig. 2.7.2 The Grotto, a cave eroded along a plane of weakness in the limestone of the Niagara escarpment in Bruce Peninsula National Park, Lake Huron.

⊡⊡ Fig. 2.7.3 Singing Sands beach on the Bruce Peninsula, Lake Huron. A thin veneer of sand overlies limestone.

of 592 km on the mainland and 47 km on islands. Lake Erie, the smallest in volume and shallowest of the Great Lakes, has a coastline dominated by outcrops of glacial till. The lake can be divided into three distinct basins. The western basin is the shallowest and the north coast is composed of 3–20-m high bluffs of glacial drift that decline to sediment-starved beaches

(>Fig. 2.7.6). This basin extends from the Detroit River to Point Pelee, the southernmost part of mainland Canada. Recently, under relatively low lake levels, erosion has been a problem in this part of the lake where the lakebed is composed of cohesive clay, and new models are needed to predict the outcomes (Trenhaile 2008). The central

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⊡⊡ Fig. 2.7.4 Vegetated sand dunes at Pinery Provincial Park, Lake Huron, formed as lake level fell during the past 10,000 years. The sand deposits extend at least a kilometre inland.

⊡⊡ Fig. 2.7.5 The beach at Pinery Provincial Park, Canadian coast Lake Huron, in winter. Often the beach is frozen, with several ice ridges developing offshore. These contain an abundance of sand that melts out in places in spring, and is added to the beach.

basin, extending from Point Pelee to Longpoint has a north coast composed of bluffs of glacial till up to 40-m high (>Fig. 2.7.7). The bluffs are easily eroded and the nearshore waters are often heavily loaded with sediment (dominantly silt and clay). On the eastern end of this

stretch, the bluffs are capped by the sand of the Norfolk sand plain, and cliff-top sand dunes are common. The eastern basin extends from Longpoint eastward to the Niagara River. The north coast of this basin is composed of limestone outcrops and up to 15-m high glacial bluffs.

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⊡⊡ Fig. 2.7.6 Beach sand deficit on the NW coast of Lake Erie has prompted coast protection works. Block armouring has been introduced just east of Point Pelee, with a groyne that has failed to intercept longshore drifting sediment.

⊡⊡ Fig. 2.7.7 Bluffs on the coast of Lake Erie, cut in soft glacial till and capped by thin deposits of sand, forming cliff-top dunes.

Pocket beaches are common between the limestone headlands. Lake Erie empties into the Niagara River and Welland Canal, which open into Lake Ontario.

5.  Lake Ontario Lake Ontario (74 m above sea level) is the smallest lake in surface area. However, its volume is large due to its great

depth, the average being over 86-m deep and the deepest part more than 244 m below low water datum. Depths increase rapidly offshore in the west, while the east sees a more gradual deepening. The Canadian coastline of Lake Ontario is more than 537 km long, with an additional 80 km on islands. The coastline of the western end of the lake is comprised of bluffs of glacial till and outwash. These are up to 90-m high in the Scarborough bluffs in the east of Metro Toronto (Eyles 2002). The bluffs decrease in size to

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⊡⊡ Fig. 2.7.8 Gently sloping beach and small dunes at Presquile Provincial Park, Lake Ontario. Vegetation has been planted to trap wind-blown sand and prevent it blowing inland.

the east where marshes and barrier dune complexes have developed. Presquile and Sandbanks Provincial Parks are typical of barrier dune and marsh complexes (>Fig. 2.7.8). Prince Edward County is a limestone plateau at the eastern end of the lake with an irregular coast containing bays that are partly blocked by spits or barrier islands.

References Carter CH, Haras WS (1985) Great lakes. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 253–260

Dech JP, Maun MA, Pazner MI (2005) Blowout dynamics on Lake Huron sand dunes: analysis of digital multispectral data from colour air photos. Catena 60:165–180 Eyles N (2002) Ontario rocks. Fitzhenry and Whiteside, Markham, Ontario, p 399 Great Lakes Atlas (2000) The Great Lakes: An environmental Atlas and Resource. Book http://www.on.ec.gc.ca/great-lakes-atlas/intro.html Hough JL (1958) Geology of the Great Lakes. University of Illinois Press, Urbana, IL Houser C, Greenwood B (2005) Profile response of a lacustrine multiple barred nearshore to a sequence of storm events. Geomorphology 69:118–137 Trenhaile AS (2008) Modeling the erosion of cohesive clay coasts. Coastal Eng 56:59–72

3.0 Pacific Central America – Editorial Introduction

Central America is here taken to extend from Mexico south to Colombia. It is dominated by mountain ranges that continue the trend of the Rocky Mountains of North America southward through an isthmus to the Andes Mountains of South America, but there are intricate structures resulting from a complex tectonic history, and some associated vulcanicity. Rock formations range from Palae­ ozoic to Holocene (Weyl 1980; Yanez-Arancibia 2005).

1.  Pacific Coast On the Pacific coast narrow coastal plains alternate with steep sectors. Earthquakes occur, and there are many uplifted beaches and terraces on the coastal slopes. Ocean swell from the SW breaks on beaches supplied with sand mainly from rivers draining the steep slopes of coastal ranges. Beaches pass laterally into barriers in front of coastal lagoons and swamps. In Costa Rica and Panama the coastline becomes irregular in response to geological structures, with long sectors of steep and rocky coast, and shallow embayments, notably the Gulf of Panama, in which wave action is attenuated. Mean spring tide range is generally small (Fig.  3.2.1). This is about 2,200 km overland. There is considerable geological and geomorphological diversity from one end to the other (>Fig.  3.2.2) (Yanez-Arancibia 2005). There are no known Pre-Cambrian rocks in Central America, the oldest formations being the Maya Series of  Palaeozoic metasediment in Belize (Moody 1975). The Guatemala Massif metamorphics are also of Palaeo­ zoic age. The Mesozoic is represented by Triassic and Jurassic metamorphic rocks in the Honduras Massif, and on the Santa Helena peninsulaby Triassic and Jurassic dark shales near Tegucigalpa and Todos, the Santos red  beds in Guatemala, and by a variety of Cretaceous sediment and volcanic formations. Cenozoic volcanic and sedimentary rocks are widespread, and include Palaeocene sediment in the Peten lowland, the thick Eocene clastic beds of the Toledo Formation in Guatemala and Eocene volcanics in Costa Rica; Oligocene and Miocene clastic sediment, limestones, and volcanic rocks in Costa Rica, Nicaragua and Panama, and marine Pliocene formations in Panama. The most recent volcanic rocks in the region are of Pleistocene and Holocene age (Weyl 1980). There are three geotectonic provinces in Central America: (1) the Peten lowland, a continuation of the Campeche-Yucatan basin, in Guatemala and Belize; (2) the fault block mountainous massifs in Belize, Guate­­ mala and Honduras; and (3) the volcanic terrain made up  of the Isthmian link – the Pacific coastal plain, the Nicaraguan volcanic upland and the volcanic ranges of the Pacific ­cordillera from Mexico to Panama. The generally accepted plate tectonic theory concerning Central America holds that the region has been thrust northeastward, as

part of the Caribbean-East Pacific plate, between the essentially westward moving North and South American plates. Tides on the Pacific coast of Central America are semi-diurnal. The tide range is 1.5–2 m at Guatemala, 2–3 m at El Salvador and Nicaragua, 4 m in the Gulf of Fonseca and 4.5–6 m in the Gulf of Panama (GierloffEmden 1982). The coast is exposed to Pacific ocean swell, with waves in the surf zone up to 2 m high (GierloffEmden 1982). The predominant longshore drifting is southeastward.

2.  Guatemala The Pacific coast of Guatemala is about 250 km long and consists essentially of sandy beaches, beach ridges and dunes, with no cliffs. The coastline is gently curving because sand supplied by longshore drift has sealed former inlets. Some are now lagoons, bordered by mangroves or bare hypersaline tidal flats. River mouths have been deflected behind beach ridges, and deltas are blunted by high wave energy over a narrow continental shelf. Beach and dune biota include the creeping vines Ipomea and Canavalia, erect plants like Uniola, Scaevola and Distichlis, and burrowing invertebrates such as lamellibranchs, polychaetes and crustaceans.

3.  El Savador The northwest coast of El Salvador is rocky and cliffy, with red and green algae in the sublittoral zone. Barnacles, chitons, crabs, mussels, periwinkles and limpets are found in the tide pools and on the rocks. Barrier islands dominate the central sector of the coast (>Fig. 3.2.3). The barriers are 1–2 m high and the lagoons behind them are 1–3 m deep, with tidal channels to a depth of about 10 m, and a

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Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

⊡⊡ Fig. 3.2.1 Location map of the Pacific coast of Central America, showing drainage patterns. (Courtesy Geostudies.)

dense mangrove thicket with meandering tidal channels. The southeast coast of El Salvador, outside the Gulf of Fonseca, has sandy beaches backed by dunes. The Gulf of Fonseca is flanked on the west­­ern  side by the high Conchagua Volcano (1,258 m), and has  a mangrovefringed northern coast (Gierloff-Emden 1959).

4.  Honduras The Pacific coast of Honduras is a 70 km long deltaic sector behind the Gulf of Fonseca. This Gulf, bordered by El Salvador, Honduras and Nicaragua, extends about 50 km inland and is 70 km wide, backed by a largely mangrove fringed deltaic coastline. The largest river entering the gulf is  the Goascoran, on the border between El Salvador and

Honduras. There are three volcanic islands, each 5–10 km in diameter, in the gulf: El Tigre, Conchaquita and Meanguera.

5.  Nicaragua Southeast of the Gulf of Fonseca, the coast of Nicaragua, more than 300 km long, includes the mountainous ­volcanic Cosiguina Peninsula. To the southeast extensive beaches and beach ridges are cut by small mangrove-fringed rivers. Longshore drifting is from south to north. At the southeastern end the Nicaraguan coast becomes steep with rocky shores and coves with pocket beaches separated by cliffs. On 2 September 1992 an earthquake SW of Managua caused a tsunami of up to 10.7 m along about 200 km of the coastline.

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

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⊡⊡ Fig. 3.2.2 Geomorphological map of the Pacific coast of Central America. (Courtesy Geostudies.)

6.  Costa Rica

7.  Panama

Steep and rocky coasts extend along most of Costa Rica. Plant and animal communities are like those of northern EI Salvador. Emerged terraces are prominent along some stretches of this steep coast. The peninsula of Capo Sta. Elena follows an east–west anticlinal trend, and has drowned valleys along its northern coast. The Nicoya peninsula has marine terraces deformed by tectonic movements, and has been subject to many tsunamis (Fernandez et al. 2000). The Gulf of Nicoya is a wide indentation bordered by sandy beaches, with a northwest extension to the mouth of a mangrove fringed estuary. Longshore drifting from the southeast and east has fed the long spit of Puntarenas, backed by a dense mangrove forest. The large open Bahia de Coronado shows blunt deltas and barred river mouths with beach ridges and mangroves in swales on coastal barriers. To the south east cliffs and bays with small beaches border the Peninsula de Osa and the Golfo Dulce. The gulfs occupy grabens between uplifted mountain chains.

Panama has the longest Pacific coastline (more than 900 km) of Central America. Indented rocky shores with steep bluffs border the coast of the Gulf of Chiriqui, with the high outlying volcanic island of Coiba rising to 400 m and the steeply-sided Peninsula of Azuero to the east. To the east the steep coasts are interrupted by bays with barrier islands, beach ridges, blunt deltas and mangroves. High rainfall feeds rivers draining steep hinterlands, which have delivered large quantities of sediment to the coast. The Gulf of Panama is about 180 km long from east to west and 140 km from north to south, and is rather shallow with depths of only 50–100 m in the central region. The western and eastern coasts are steep and rocky but intertidal flats are extensive along the coast west of the Panama Canal, and to the east are large areas of marsh and mangrove swamps. The indented coastline to the east includes submerged valley mouths bordering parallel mountain chains. The main coral reefs on the Pacific coast of Central America are located in Panama off the Azuera

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⊡⊡ Fig. 3.2.3 The coast of El Salvador between the Rio Lempa and the Rio Grande de San Miguel, showing beach ridge patterns on the outer barrier and a segment of an earlier inner barrier. Mangrove swamps are shown in green. (Courtesy H.G. Gierloff-Emden.)

190 Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama

Peninsula and Coiba Island, in clear water away from river mouths and suspended sediment.

References Fernandez M, Molina E, Havskov J, Aatakan K (2000) Tsunamis and tsunami hazards in Central America. Nat Hazards 22:91–116 Gierloff-Emden HG (1959) Die Küste von El Salvador. Dtsch Hydrogr Z 12:14–24

3.2

Gierloff-Emden HG (1982) Central America, coastal morphology. In: Schwartz ML (ed) The Encyclopedia of Beaches and Coastal Environments. Hutchinson and Ross, Stroudsburg, PA, pp 188–191 Moody JD (1975) Central America regional review. In: Fairbridge RW (ed) The Encyclopedia of World Regional Geology, part 2. Dowden, Hutchinson and Ross, Stroudsburg, PA, pp 228–237 Weyl R (1980) Geology of Central America. Bornträger, Berlin Yanez-Arancibia A (2005) Middle America, coastal ecology and geomorphology. In: Schwartz M (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 639–645

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3.3 Pacific Coast of Colombia

Iván Correa · Robert Morton

1.  Introduction The Pacific coast of Colombia (for its Caribbean Coast, > Caribbean Coast of Colombia) extends to the west of the Cordillera Occidental, the westernmost branch of the South American Andes. The region, known for its heavy rainfall, numerous rivers, and luxuriant vegetation, remains mostly uninhabited and undeveloped. No roads reach this coast, so access is only by boat. Overland access from the interior of the country is only possible to the cities of Bahía Solano, Buenaventura, Guapi and Tumaco, the main commerce centres of the region. Because the Nazca plate is being subducted under the South American plate, the Pacific Coast of Colombia is an active seismic zone with a well-documented record of high-magnitude earthquakes. Seismic observations for the past century record at least three major seismic events (1906, 1979 and 1992), accompanied by subsidence, soil liquefaction, and extensive landsliding that caused abundant river damming. The 1906 and 1979 events caused littoral co-seismic subsidence up to 1.5 m along the southern Pacific Coast between Buenaventura and the Mira delta. Both events triggered tsunami waves 2.5 m high that caused more than 3,500 deaths of inhabitants of the low coastal areas (Herd et al. 1981). Between Punta Ardita (at the Panamá border) and Cabo Corrientes the coastal zone corresponds to the western flanks of the Baudó Range (Serranía del Baudó), a 600  m high region composed predominantly of oceanic basalts, diabases, and associated cherts and radiolarites. The Baudó Range is a structurally controlled relief, densely fractured and faulted in the preferential directions NNESSE and N 60º E – N 30º W that control the general orientation of valleys and ridges. South of the Baudó Range, the relief of the Pacific Coast is dominated by 20–100 m high hills cut into Tertiary sedimentary sequences and Plio-Quaternary deposits at the piedmont of the Cordillera Occidental (Martínez et al. 1995). Conspicuous terraces between 20 and 30 m above present sea level are found along the hills between Cabo Corrientes and Buenaventura Bay.

Coastal relief south of the Baudó Range abruptly limits the Holocene prisms of accretionary deposits that characterise the Pacific Coast of Colombia. They include the three major Plio-Quaternary deltaic prisms of the San Juan, Patía, and Mira Rivers, and two narrow (Fig.  3.3.1), at the southern part of the delta, shows a general pattern of barrier narrowing and breaching that is becoming common along other barrier islands of the Pacific coast, including Charambirá (northern part of the delta), and along the Patía delta and Mira deltas to the south. According to radiocarbon dates, the Isla El Choncho barrier island began forming about 500 year bp., and grew to the southeast by successive beach-ridge accretion (González and Correa 2001). Its evolution in the last 50 years records a succession of events that segmented the island. The general retreat of the central part of the barrier coincided with the formation of an extensive sandy tidal flat at its northern end (Boca Chavica), which diminished sand supply from the north. Shore erosion was accelerated after a 20–30 cm co-seismic subsidence event generated by the November 19, 1991 earthquake. After the 1991 subsidence, the barrier flooded from 2 to 17 times each year ­during highest spring tides. Final breaching of the barrier occurred in 1997 and coincided with several overwash events associated with positive sea level anomalies of 30 cm caused by El Niño (Morton et al. 2000). Following a prudent strategy, El Choncho village was relocated to the Santa Bárbara beaches, a beach ridge remnant located 200  m inland from the coastline (Correa and Gonzalez 2000). To the south of the San Juan delta, the coastline of Málaga Bay and the northwest side of Buenaventura Bay are characterised by vertical to sub-vertical cliffs, 10–20 m high, cut into horizontal to sub-horizontal Tertiary sandstones and mudstones. The coastal relief is characterised by low hills and a terrace level of 20 m. Erosional segments of cliff shore alternate with small pocket beaches supplied with sand produced by rapid shore retreat. The numerous stacks and arches are evidence of the retreat. Depositional zones are represented by mangrove-vegetated tidal flats at the eastern sides of Málaga and Buenaventura Bays, and by the Juan Chaco beaches (>Fig.  3.3.2), an unstable supratidal feature formed with sand from the San Juan delta. South of Málaga Bay, the coast is dominated by steep,

Pacific Coast of Colombia

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⊡⊡ Fig. 3.3.1 Central part of El Choncho barrier island after breaching of the new inlet, near El Choncho village.

plunging cliffs, 10–30 m high, cut into horizontal to ­sub-horizontal Tertiary mudstones and sandstones, with notches and evidence of bioerosion at the intertidal levels.

3.  Southern Pacific Coast The Southern Pacific coast has an overall length of 650 km between Buenaventura Bay and the Mira River delta. The coast from Buenaventura Bay to Guapi Bay extends for about 150 km, with a general S 40º W trend. The southern Buenaventura Bay shore is supplied by the Anchicayá and Dagua Rivers and is a low depositional area characterised by segmented sand bars and extensive, intertidal mudflats and mangrove swamps that penetrate more than 10 km inland from the coast. Farther south, the littoral fringe represents a system of barrier islands mangrove and transitional swamps (>Fig.  3.3.3), and is supplied by numerous rivers. The largest are the Cajambre, Naya, and Micay Rivers that drain the western slopes of the Cordillera Occidental. A dense network of tidal channels interconnects the estuarine lagoons and mangrove interdistributary areas, and the freshwater swamps are backed by cliffs or gently sloping platforms cut into Tertiary sedimentary rocks. Modern barrier islands of this stretch are 2.8–5.6 km long and 0.2 –1.1 km wide and are mostly transgressive features (Martínez et al. 1995). The Patía River drains about 23,000 sq km of the Central and Western Cordilleras of the Colombian Andes. Its delta consists of a northern ancient delta plain and the

western, more recent delta plain. This zonation results from the progressive NE to SW delta migration. Both delta zones have been influenced by Quaternary tectonic movements. The Patía delta coastline is about 120 km long and extends from the Guapi Bay to Punta Cascajal, at the northwestern extreme of Tumaco Bay. The northern Patía delta fringe is a subsiding, tide-dominated area characterised by frontal, erosional lobes of sandy and muddy tidal flats backed by mangrove swamps and an ancient complex of beach ridges. Numerous funnel-shaped estuaries and tidal channels penetrate more than 15 km to the interior of the northwestern Patía delta plain, where mangrove and transitional swamps are rapidly eroding. The erosion produces a pattern of numerous isolated small islands. The western and southern shores of the Patía delta consist of external sets of barrier islands and mangrove and transitional swamps backed by intricate sets of point bars and sandy fluvial deposits of numerous abandoned courses of the Patía River. Barrier islands of this sector have the best dune development (2 m high) of the barrier islands of the Pacific Coast of Colombia. Nevertheless, the barriers exhibit transgressive characteristics with abundant breaches and island segments (Martínez et al. 1995). The entire Patía delta has been affected in historical times by co-seismic subsidence, best recorded in the December 12, 1979 earthquake that affected the Southern Pacific Coast from the Mira River delta to Guapi Bay. This event in particular caused a regional subsidence maximum of about 1.6 m along the shores of the San Juan barrier island, and its associated tsunami killed about 155 people at the

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Pacific Coast of Colombia

⊡⊡ Fig. 3.3.2 Juan Chaco village at the entrance of Málaga bay, showing the 10 m terrace and cliffed coastline with caves and structurally controlled entrances cut into subhorizontal mudstones and sandstones. (Courtesy Observatorio Sismológico del Suroccidente Colombiano, Universidad del Valle.)

⊡⊡ Fig. 3.3.3 Mangrove swamp (Rhizophora mangle) along the northern border of the Guapi estuary. The mangrove trees are up to 30 m high.

San Juan de la Costa village, located on the central part of this island (Herd et al. 1981). Tumaco Bay, located between the Patía and Mira deltas, is the largest embayment of the entire Pacific Coast of Colombia. The bay is a shallow (up to 30 m deep), semiprotected area located on an active tectonic area with a N 30º E structural trend. The Tumaco Bay northern and east coasts are dominated by intertidal flats and by 10–40 m

high, vertical to sub-vertical cliffs cut into horizontal to sub-horizontal mudstones. Cliff retreat is slow in most sectors, but rock falls and slumping are caused by seismic events and occasionally strong rain. Minor pocket beaches and wave-cut platforms can be found along the eastern cliffs of Tumaco Bay. Tidal estuaries with lobe-shaped bars and tidal flats backed by extensive mangrove swamps characterise the southeastern shores of Tumaco Bay. At its

Pacific Coast of Colombia

southwestern part, highly unstable sandy barrier and bar complexes migrating from the northwestern Mira delta front the internal shores of Tumaco Bay. Most of the city of Tumaco is located on the La Viciosa and El Morro islands, the two largest sandy barriers. Both barriers are subject to rapid cycles of accretion and erosion as intertidal bars are deposited and destroyed. Offshore sandy intertidal flats protected La Viciosa and El Morro from the direct impact of the 1902 and 1979 earthquake-generated tsunamis. The Mira River delta plain is a Plio-Quaternary accretionary prism that encompasses about 1600 sq km. The delta is limited to the east by hills and terraces of semiconsolidated sediment, and to the north by the extensive mangrove and freshwater swamps bordering southern Tumaco Bay. Morphological evidence suggests that the Mira delta prograded progressively from Tumaco Bay to the southwest. The older delta is recognised by an extensive system of arcuate beach ridges landward of the more recent deposits of sandy barrier spits and mangrove swamps. South of Tumaco Bay the NW Mira delta shore is wave-dominated with narrow, erosive barrier islands backed by mangrove and transitional swamps (Martínez et al. 1995). North­ eastern barrier islands and sandy intertidal shoals and bars migrate continuously to the southwestern part of Tumaco Bay. These features represent the most active sectors of the entire Colombian Pacific coast.

3.3

Acknowledgement The authors would like to thank Fondo BIDCOLCIENCIAS (Programa de Ciencias del Mar), grants 1216–024–99, 1216–09–153–96 and 1216–05–16911), the U.S. Geological Survey, and Universidad EAFIT for supporting the work that made this publication possible.

References Correa ID, González JL (2000) Coastal erosion and village relocation: a Colombian case study. Ocean Coast Manag 43:51–64 González JL, Correa ID (2001) Late Holocene evidence of coseismic subsidence on the San Juan Delta, Pacific Coast of Colombia. J Coastal Res 17:459–467 Herd DG, Leslie YT, Meyer H, Arango JL, Person WJ, Mendoza C (1981) The Great, Tumaco, Colombia Earthquake of 12 December 1979. Science 211:441–445 Martínez JO, González JL, Pilkey OH, Neal WJ (1995) Tropical barrier Islands of Colombia’s Pacific Coast. J Coastal Res 11:432–453 Morton RA, González JL, López GI, Correa ID (2000) Frequent nonstorm washover of barrier islands, Pacific Coast of Colombia. J Coastal Res 16:82–87 Restrepo JD, Kjerfve B, Correa ID, González JL (2002) Morphodynamics of a high discharge tropical delta, San Juan River, Pacific Coast of Colombia. Mar Geol 192:355–381

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4.0 South America – Editorial Introduction

South America consists of a major western mountain range, the Andes Cordillera, extending from Colombia southward through Ecuador and Peru to Chile and Patagonia, the broad plateau of southern Brazil, the uplands of Venezuela and the large river basins of the Amazon and the Parana (Psuty et al. 2005). The Andes are primarily the outcome of Alpine folding, while the plateaux of southern Brazil and the Venezuelan uplands are Pre-Cambrian shields.

1.  Pacific Coast In > Ecuador and northern > Peru the Andes mountains are bordered by coastal lowlands with hilly areas and deltaic plains. There are cliffs along hilly areas and bays excavated in less resistant outcrops. The coast is lined by beaches and dunes, and there are mangroves in estuaries and inlets. Further south the mountains run closer to the coast and there are many steep and cliffed sectors, with only minor, narrow coastal plains, generally at the mouths of river valleys. Earthquakes occur, and emerged beaches and shore terraces at various levels indicate a history of tectonic uplift and deformation. These landforms continue south into > Chile. In southern Chile the mountains are bordered by foothills, and coastal plains are longer and wider. Further south the coast becomes irregular, with many islands, and rias and (south of latitude 43 degrees) fiords dissect the mountain slopes. Contrasts result from climatic transitions from humid tropical Ecuador through arid Peru and northern Chile to cool temperate rainy areas in southern Chile, grading to subarctic in Patagonia. The coast bears the imprint of Pleistocene glaciations, and there are still glaciers in the mountains of southern Chile, several descending into the heads of fiords: in recent decades they have been receding. The continental shelf is narrow, and ocean swell from the SW dominates the ­coastline, with stormier conditions in southern Chile, where westerly gales are frequent. Man­groves extend south into Peru, with salt marshes in ­estuaries and inlets further south Coral reefs are poorly developed. Tide ranges are

generally less than 2 m, rising to 3.6 m in the Gulf of Guayaquil in Ecuador.

2.  Atlantic Coast In southern Argentina (>Argentina) the climate is cold. The hinterland was subject to Pleistocene glaciations. Beaches become gravelly, fed by rivers delivering fluvioglacial deposits and cliffs cut in moraines. In Tierra del Fuego there are frost-shattered cliffs, fiords, and gravel beaches derived from glacial drift. Ocean swell is accompanied by frequent storm waves, and tide ranges increase to more than 5 m, reaching 10.4 m at Puerto Gallegos. Further north, the pampas hills come to a coast with lengthy cliffs cut in Tertiary and older formations, interrupted by the Rio Colorado and Rio Negro deltas. Terraces at various levels are the outcome of tectonic uplift and sea level changes. South-easterly ocean swell has shaped sandy beaches backed by beach ridges and dunes. In northern Argentina the coast curves past the beachfringed lowlands of Buenos Aires on the southern side of the Gulf of the Rio de la Plata. The broad interior lowland drained by the Uruguay and Parana rivers opens to the Rio de la Plata through deltaic marshland. A hilly hinterland comes to the coast in > Uruguay, where headlands mostly contain granite, separate bay beaches backed by dunes and low cliffs form the northern shore of the Rio de la Plata. The coastal lowlands of southern > Brazil are fringed by long sandy beaches and barriers enclosing lagoons, notably the large Lagoa dos Patos. Lagoons and estuaries are fringed by salt marshes, and to the north of Cananeia there are also mangroves as the coastal climate becomes warmer. The coast north to Rio de Janeiro is hilly, with many irregular headlands, islands, embayments and some lagoons enclosed by sandy barriers. It is exposed to a strong SE swell from the Southern Atlantic. North of Cabo Frio there are several deltas, of which the Rio Sao Francisco is the most substantial, and the coastline is bordered by sandy beaches and biogenic reefs. The great crystalline Brazilian plateau ends in an

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escarpment that overlooks a coastal plain of varying width, fed by many rivers and interrupted by ridges that culminate in cliffy headlands. The swell from the Southern Atlantic weakens northward, particularly beyond Recife. The climate becomes subtropical as the coast swings north-westward. The vast Amazon delta then forms a long, low coastline with estuaries and rias, mangroves and occasional sandy beaches and cheniers. The Pre-Cambrian uplands of Bolivar do not extend to the coast, which consists of the wide alluvial plains of > French Guiana, > Surinam and > Guyana. Mean spring tide ranges diminish to less than 2 m, and the coast is exposed to NE waves from the Atlantic Ocean.

3.  Western Caribbean Coast In eastern > Venezuela the Orinoco valley opens to a major delta. The climate is humid tropical, with occasional

hurricanes. Dominant waves are from the NE and mean spring tide ranges are small ( Fig. 4.1.1). Sandy beaches and sloping cliffs continue south past Punta Charapato to Bahia de Manta, where the coast curves westward past Manta to a bold promontory of PlioQuaternary sediment over Miocene mudrock. Manta is a seaport with a harbour protected by breakwaters built in 1972 and later extended, and these have intercepted sand that drifts alongshore during northwesterly swells to form a wide beach on the western side. The narrow, depleted beach to the east led to accelerated cliff recession, which was countered by the building of a sea wall. At Tarqui there has been pollution resulting from artisanal fish cleaning and gutting on the beach in a sector where cleansing wave and current scour has been diminished by nearby harbour structures. Sloping cliffs in soft silty clay run westward behind Playa Murcielego (> Fig. 4.1.2), their recession threatening cliff-top houses (> Fig. 4.1.3). The sandy beaches are highly calcareous and often shelly, and quartz is the most abundant terrigenous component. The cliffed coast swings southward to Cabo San Lorenzo, where Cretaceous volcanic rocks rise to outcrop

⊡⊡ Fig. 4.1.1 Beach in the south of Bahia de Caraquez. (Courtesy Geostudies.)

Ecuador

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⊡⊡ Fig. 4.1.2 Cliffed coast west of Manta. (Courtesy Geostudies.)

⊡⊡ Fig. 4.1.3 Walled cliff sector near Manta. (Courtesy Geostudies.)

on the shore and there are cobble beaches. Upper Cretaceous sediment form hilly cliffed promontories on the coast past Punta Cayo (> Fig. 4.1.4) and there are sandy valley-mouth plains at Machalilla and Puerta Lopez, with beaches shaped by southwesterly ocean swell refracted to approach from the west.

The beaches at Machalilla and Puerto Lopez include sparse coral gravel, swept in from patchy reef gardens on rocky foundations in shallow nearshore areas. Lower Cretaceous volcanic rocks outcrop in the Colonche Mountains, just inland. Regional uplift is indicated by the occurrence of tablazos which have been mapped as far

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as 27 km inland. There is a sharp cliffy headland at Punta Lopez (> Fig. 4.1.5), where steep-sided Salanco Island stands offshore. The coastal region is dry, with areas of semi-arid scrubland. The beaches are of fine grey calcareous sand, with shelly patches and gravels near rocky outcrops in cliffs and headlands.

The sandy beach widens southward past Salanco, a resort town, and the mouth of Rio Ayambe to Mantañita Point. There are rocky stacks, encrusted with seabird guano. Manglaralto and Valdivia are in bays with sandy beaches receiving surf from southwesterly swell, and river mouths, such as the Rio Cadeate, are often blocked ⊡⊡ Fig. 4.1.4 Cliffs south of Puerto Cayo. (Courtesy Geostudies.)

⊡⊡ Fig. 4.1.5 Cliffed headland at Punta Lopez. (Courtesy Geostudies.)

Ecuador

by swash-built sand banks. To the south the sloping cliffs in soft silty clays resume. Flat-topped ridges run out to coastal headlands between bays where the sandy beaches are backed by beach ridge plains formed by intermittent progradation of the sandy coastline. There are segments of shore platform cut in indurated mudrock on the headlands. In places the coast has been cut back, threatening the coast road at San Pablo, and erosion of the beach has exposed beach rock, broken and piled up by storm waves at Punta Baruynda, to the south (> Fig. 4.1.6). The coast curves round the Bahia de Santa Elena to Salinas, a seaside town on a dry peninsula, with an esplanade behind a broad sandy beach sheltered from southwesterly swell but receiving northwesterly swell between November and March. The tabular peninsula of La Puntilla is cut in Eocene sandstone underlain by Upper Cretaceous silts and agglomerates. The coastline swings southeastward, and is thus directly exposed to southwesterly ocean swell. The relatively straight coastline crosses a hilly area, with sectors of cliff cut in MioPliocene and Quaternary sediment, some of which show landslides. These alternate with valley-mouth lowlands, where dunes have formed behind the beach, as between Chanduy and Point Piedras. The broad beaches of calcareous sand contain deposits of dark magnetite and titanium sands, which have been mined. Ocean swell breaks across sand bars in the nearshore zone, and river mouths are often dammed by barriers to form small lagoons. The coastal

⊡⊡ Fig. 4.1.6 Disintegrated beach rock south of San Pablo. (Courtesy Geostudies.)

4.1

region is dry enough for salt making, and salt water intake structures interrupt the beach west of Punta Carnero. Playas is a large seaside resort with a wide beach, popular with people from Guayaquil, and at Posorja a spit has formed as the coast curves into the Gulf of Guayaquil. The Gulf of Guayaquil is a triangular-shaped area submerged during the Late Quaternary marine transgression and partly filled with sediment brought down by the Guayas and other smaller rivers (Ayon 1976). It contains Isla Puna, a large island of Plio-Quaternary sediment (including tablazos) overlying Miocene mudrocks, which separates a narrow western entrance, Canal del Morro and a broader eastern entrance, Canal de Jambeli. There are a few low cliffs and shore platforms on the more exposed western coast, but much of the island is fringed by sandy beaches and spits enclosing tidal mudflats and mangrovefringed lagoons. Mangroves and salt marshes occupy the concave eastern coast, and there are recurved spits at the southwestern end of the island. The Canal de Morro and Canal de Jambeli lead north into branching mangrove-fringed inlets and swampy deltaic islands, and the Guayas estuary, beside which stands the city of Guayaquil. There is much floating Eichhornia, derived from the freshwater swamps north of Guayaquil. Tidal currents are strong, and there is an extensive and variable ebband-flood shoal and channel topography with megaripples. On the western side is the intricate Estero Salado, a channel about 42 km long in which tides attain 5 m. As mangroves

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Ecuador

spread on to accreting mudflats intervening channels are scoured more deeply (Cintron 1979). On the eastern side of the Gulf of Guayaquil a swampy coastal lowland extends south to Machala. Small intertidal deltas emerge during low tide at the mouths of the rivers running at right angles to the shore through the mangrove fringe. Shrimp (prawn) ponds have been excavated in salt marshes and saline flats, and in the mangroves. Emerged deposits of Crassostrea shells, up to 10 m above sea level, are evidence of Holocene uplift. The coastal plain ends in the mangrove-fringed Jambeli Archipelago which extends over an area of 260 sq. km, close to the Peruvian border. Sandy beaches develop as wave energy increases south of the protection of Puna Island, and sand spits are growing eastward from river mouths on the Tumbes delta.

3.  Galapagos Islands These islands form the Galapagos Province of Ecuador, and are famous for their unusual fauna, including giant turtles, described by Charles Darwin when he visited the islands in 1835 during the voyage of the Beagle. They are a group of 19 volcanic islands (and numerous islets and reefs) on the Equator in the eastern Pacific Ocean, about 965 km west of the coast of Ecuador. They formed as shield volcanoes rising 3,000 m from ocean floor fractures, and consist largely of superimposed basalt lava flows. They formed in response to a hot spot that has migrated westward, so that the older islands with extinct volcanoes are in the eastern part and the younger ones with active volcanoes to the west. Isla Fernandina is a volcanic cone rising to La Cumbre (1,463 m). It occupies a bay on the western side of Isla Isabela, a large curved island with several volcanic

peaks, the highest of which is Cerra Azul (1,689 m) in the north. Isla San Salvador, Isla Pinzon, Isla Santa Cruz, Isla Santa Maria and Isla Espando are all simple volcanic islands. Eruptions have continued on a small scale. The occurrence of submarine lavas and Pliocene sedimentary rocks indicate that some of the islands have been raised about 100 m above sea level. The climate is dry and cool as a result of the cold Humboldt Current arriving from the south, and the islands are sparsely vegetated, with extensive rocky areas. Wave action is stronger on the south and west facing coasts, which receive swell transmitted from the Southern Ocean across the South Pacific. The coasts of the volcanic Galapagos Islands show rugged cliffs, shore platforms, some dark beaches of basaltic sand and gravel and a few beaches of white calcareous (coralline and shelly) sand, notably on the older eastern islands. Perry Isthmus on Isla Isabela is a sandy tombolo. Some of the steeply plunging cliffs are associated with laterally eroded craters and calderas. Dissection along joints and planes between lava flows has guided the shaping of caves and natural arches. Shore platforms are developed on seaward protruding lava flows. There is a large blowhole at the western end of Isla Espando, and one of the islets is pierced by a natural arch.

References Ayon H (1976) Sediments of the Guayaquil estuary. Instituto Ocean­ ografico de la Armada del Ecuador Bird ECF (1983) Shore procersses in Ecuador. In: Vallejo S (ed) Ordenacion y Desarrollo Integral de las Zonas Costeras, Armada del Ecuador, pp 29–45 Cintron G (1979) Mangroves of the Ecuador coast. Departamento de Recursos Naturales, Puerto Rico

4.2 Peru

Vidal Taype Ramos · Eric Bird

1.  Introduction The coastline of Peru (>Fig. 4.2.1), just over 2,300 km long, is notably arid (Dresch 1961). Mean annual rainfall at coastal centres is low (Chiclayo 15 mm, Trujillo 28 mm, Lima 31 mm and Mollendo 22 mm) and variable; desiccation is intense between the occasional rains. This is because coast­­al waters are abnormally cool for the latitude (4°–18° S), due to the upwelling of cold water from oceanic depths as a result of trade winds blowing offshore. Sea breezes keep the coastal fringe cool, with frequent fogs but very little precipitation. Vegetation is consequently very sparse, and desert landscapes predominate. Tectonic uplift accompanied by faulting has raised the high hinterland of the Andes Mountains, the structural grain of which is followed by the coastline (Kellog and Mohriak 2001). Coast Ranges, extensive in Chile, are here confined to the sector between Pisco and Mollendo (Noble et al. 1978) with an outlier near Punta Aguja; intervening segments have subsided beneath the sea. Earthquakes still are found. In May 1970 a major earthquake centred 50 km off Chimbote caused massive avalanches in the mountains and severe flooding down the Santa River valley to its delta region. Many valleys are incised into the coastal slopes, but only streams fed by mountain rainfall and melting snow maintain their flow through to the coast, and even these show great variations in seasonal and annual discharge. The coastline is relatively simple in outline, with few estuarine inlets and only minor deltaic protrusions built in the face of oceanic wave action (swell generated in the Southern Ocean crosses the South Pacific to arrive on this coast from the south-west. Cliffs cut by marine erosion alternate with sandy beaches derived mainly from shoreward drifting of sea-floor sediment and some sand and shingle beaches nourished from cliff erosion and fluvial outwash. Beaches are backed by unvegetated dunes, or have only a sparse plant cover sustained by fog and dew (Dresch 1961). In many places sand is drifting inland to be banked against cliffs or rising slopes, or to block the mouths of coastal valleys and gullies. Ocean swell arriving from the southwest is often accompanied by small waves generated in coastal waters by southerly breezes. These waves arrive at an angle to the

coastline and cause northward drifting along beaches. Occasionally in summer there are northwesterly waves (which cause southward drifting) and penetration by the warm El Nino current to the coast north of Negritos. Mean spring tide ranges are small, generally less than a metre (0.7 m at Callao, the port of Lima), but they in­­ crease northward to the Ecuador border and the Gulf of Guayaquil.

2.  The Peruvian Coastline Close to the Ecuador border is the large, sandy Tumbes Delta on the southern shores of the Gulf of Guayaquil. Its estuarine distributaries are mangrove-fringed (the southern limit of mangroves on the west coast of South America is here at Punta Malpelo, 3° 4' S, because northward drifting carries seedlings in that direction). Parallel sandy beach, ridges, emphasised by lines of shrubs, extend southward to Caleta Le Cruz, the landing place of Fransisco Pizarro in 1527. Southwest from Zorritos the coastline is generally cliffed, cut into a fringe of uplifted Tertiary formations that dip gently seaward and are capped by a series of Quaternary marine terraces (tablazos), incised by transverse valleys (quebradas), but otherwise well preserved in this arid climate. Each terrace is a sparsely vegetated plain strewn with broken, sand-blasted pebbles, with old beach deposits to the rear, beneath bluffs that mark degraded cliffy coastlines. The oldest and highest, the Mancora terrace, stands about 360 m above sea level in the north and declines to 75 m southward. The next, the Talara terrace, falls southward from 90 to 40 m, while the youngest and lowest, the Lobitos terrace, declines from 40 to 30 m southward (Bosworth 1922). These terraces are truncated by cliffs along the present coastline: at Cabo Blanco, for example, the Mancora terrace ends in cliffs 300 m high. An early Holocene coastline is marked by gullied bluffs behind intervening sectors of coastal plain, as at Talara. The beaches are of shelly sand with some stony material, and the coastal plains include parallel beach ridges (Wells 1996), often overlain by dunes that have locally been banked against former cliffs 45–60 m high.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4. 2

Peru

⊡⊡ Fig. 4.2.1 The Peruvian coastline. (Courtesy Geostudies.)

THE PERUVIAN COASTLINE

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Near Negritos the cliffs are 80 m high, incised by valley mouths blocked by 20 m high dunes. Southeast from Parinas Point the sandy beach, backed by dunes, is 50 km long, interrupted only by the mouth of the Chira River. Behind the dunes is a flat plain inundated by the highest tides, and covered with evaporite deposits.

75°

Tam

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70°

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208

Paita is on a hilly peninsula of Mesozoic rock, but the bay to the south is sandy, fronting abandoned cliffs, and paired spits border the small estuary of the Rio Piura. Punta Agula is another hilly peninsula dissected by wadis, and followed southward by a sandy coastline fringing the Sechura Desert. The extensive dunes here include

Peru

4. 2

⊡⊡ Fig. 4.2.2 Indented coastline north of Huarmey. (Courtesy Geostudies.)

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barchans up to 6 m high and 15 m wide. Evaporation is intense, and the Rio Cascaja fades away before reaching the sea. The coastal plain narrows southward from Chiclayo to Trujillo, beyond which a steep coast predominates as the Andean foothills approach and intersect the coastline. Cliffs and bluffs are cut into Mesozoic rocks and intruded granites and there are rugged islands offshore (e.g., Isla de Macabi and Islas Guanape), many with thick guano deposits. Sectors of sandy beach (playa) also are found, backed by beach ridges, but south of Chimbote the coastline is hilly and embayed on Cenozoic volcanics, with shore platforms fronting cliffs and the harder elements persisting as rocky headlands, with caves and coves. The Peninsula de Ferrol is a tombolo, a high hill at­­ tached to the mainland by a sandy isthmus; its rocky seaward margin has bouldery clefts and a natural tunnel. A scalloped cliffy coastline (>Fig. 4.2.2), notched by wadis, near Huarmey, is fringed by many white or yellow guano-mantled islands and headlands. The Rio Pativilca has built a blunt delta, while to the south Les Salinas is a small elongated barrier lagoon, and Playa Grande has

heavy surf breaking on a shingle beach fronting bluffs cut into the seaward margin of a gravelly pediment. Callao, the port of Lima, has large breakwaters enclosing a harbour, with accretion of sand from Rio Rimac immediately to the north. A major spit (La Punta) protrudes in the lee of Isla San Lorenzo, a hilly offshore island on which Quaternary marine terraces are well developed. Southward the coast is undulating, with small plains at valley mouths, and the Paracas promontory at Pisco marks the beginning of a coastal mountain range. Punta Zarata is a cuspate foreland enclosing a lagoon and salt marshes, and a spit partly encloses Laguna Grande on the east coast  of the Carreta Peninsula, bordering Bahia de la Independencia. Asymmetrical bays, shaped by the refraction of southwesterly swell around headlands, have developed at Bahia de Lomitas, Bahia San Nicolas, and Bahia San Juan. Lo­­ cally, pebble beaches have formed from material derived from Pleistocene conglomerates deposited at old valley mouths and now exposed in cliff sections. The next part of the coastline faces southwest and receives strong surf produced by swell arriving through

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Peru

⊡⊡ Fig. 4.2.3 The blunt delta of the Rio Tambo, built on a coastline exposed to southwesterly ocean swell, which has built barrier beaches impounding small, narrow coastal lagoons. (Courtesy Geostudies.) 60 0

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Peru

deep water inshore. Drifting sand blocks the mouth of the Ico River valley. To the south the coast is steep on granitic rocks, with many gullies, some of which reach the sea in small coves. Rio Ocona has a small estuary, but Rio Camana has built a blunt delta with beach ridges and swale lagoons, as has the Rio Tambo southeast of Mollendo (>Fig. 4.2.3). Rio Vitor, by contrast, reaches the coast through a deep gorge. Between Mollendo and the Chilean border are several straight beaches fronting bluffs that truncate pediments: Playa de El Palo (>Fig. 4.2.4), Playa Tacahuay and Playa Inglesa.

4. 2

References Bosworth TO (1922) Geology and palaeontology of Northwest Peru. Macmillan, London Dresch J (1961) Observations sur le désert cotier du Perou. Ann Géog 70:179–184 Kellog JN, Mohriak WU (2001) The tectonic and geological environment of coastal South America. In: Seeliger U, Kjerfve B (eds) Coastal marine ecosystems of Latin America. Ecological Studies 144:1–16 Noble DC, McKee EH, Megard F (1978) Eocene uplift and unroofing of the coastal batholith near Lima, Peru. J Geol 86:403–405 Wells LE (1996) The Santa beach ridge complex: sea-level and progradational history of an open gravel coast in central Peru. J Coastal Res 12:1–17

211

4.3  Chile   José Araya-Vergara  

1.  Introduction The coast of Chile is about 6,435 km long. Tide ranges are small in the north, Arica having a mean spring tide range of 1.4 m and Iquique 1.5 m. They increase southward to Valparaiso (1.66 m) and Caleto Mansa (1.62 m), and are much larger in the Gulf of Corcovado (Puerto Montt 5.8 m). Further south they are generally between 1.2 and 1.4 m, rising to 2.06 m in Orange Bay, near Cape Horn. Wave action is dominated by the Pacific ocean swell from the southwest and storms in coastal waters that produce waves mainly from the west. The climate is hot and dry in the northern desert, mild and Mediterranean in the centre (winter rainfall and dry summers: Antofagasta has mean monthly temperatures of 20.6°C in January and 14°C in July, with an average annual rainfall of 12 mm) and cold, wet and stormy in the Patagonia to the south.

s­ upplied directly from the Copiapó River (Araya-Vergara 2001), but dune sands of the ergs located in basins many kilometres inland had a marine origin. South from the Copiapó River, all the principal rivers flow permanently to the sea through microtidal estuaries (>Fig. 4.3.4). Over more than 700 km of coastline, increasing supplies of fluvial sediment to the shore form beaches in zeta form embayments. Northward drifting within these results in wider and higher beaches and dunes at the northern ends, as in Tongoy Bay and the Los Vilos embayment. Between the estuaries of Petorca-Ligua and Maipo rivers there is a southward transition from rias to estuarine deltas, as in the Aconcagua estuary (>Fig. 4.3.5).

⊡⊡ Fig. 4.3.1 Map of northern Chile. 71°W 18°S

2.  The Chilean Coastline

Iquique

1110 km

South from the Peruvian border at Arica the coast of the Atacama desert is mainly cliffed (>Fig. 4.3.1). Near Iquique, massive transgressive dunes (>Fig. 4.3.2) have been deposited between the present beach and the local mega-cliff. Due to their position with respect to present sea level and the degree of weathering of their sands (Potter 1994) they are thought to have formed during the Pleistocene (Paskoff et al. 1998), when sea level was lower. They are a striking feature, because dunes are scarce on the cliffy coast of the desert. North of Antofagasta cliffs are cut in horizontally bedded Tertiary sandstones (>Fig. 4.3.3). In some parts the coastline has embayments cut in sandstone between protruding headlands where harder rocks outcrop near mean sea level. There are emerged terraces, as at Cobija, where a 15-m Last Interglacial terrace fronts an escarpment. At Caldera, south of Antofagasta the Copiapó River is the first river outlet with permanent runoff to the sea, after more than 1,000 km of desert coast. There are Pleistocene marine terraces with deflated beaches on their surfaces (Marquardt et al. 2000), and the ergs of the Atacama Desert. A large part of the sand of the ergs has been

Arica

Ancient transgressive dunes

La Portada: Soft cliffs and arces Antofagasta 71°W 25°S

Caldera Copiapó River

Ergs Marginal Desert of Atacama Coastal terraces isotopically correlated First river oulet with permanent runoff to the sea in the marginal desert

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

214

4. 3

Chile

⊡⊡ Fig. 4.3.2 Transgressive dunes near Iquique, fronting a megacliff. The dune crests change in relation to wind direction.

⊡⊡ Fig. 4.3.3 Sandstone cliffs at Antofagasta, showing contrast between slope on upper softer beds and cliff on harder sandstone.

Between the Maipo River outlet and San Antonio harbour the coastline has been straightened as the result of accretion alongside a breakwater. The coastline advanced 600 m between 1908 and 1934 forming beach ridges, foredunes and a lagoon, but since 1934 the ­coastline has stabilised (Araya-Vergara 1985). At Santo Domingo sand formed beach ridges and successive parallel foredunes (>Fig.  4.3.6), four of which were added during the past 600 years (Paskoff et al. 2000). The penultimate ridge was monitored during the storm surges of 1968 and 1982, when the dune was cliffed. After 1982 a wide beach formed and microdunes developed, and by 1993 the present seaward foredune had been built (>Fig. 4.3.7). South of Santo Domingo barchans behind the beach diminished in number and area between 1955 and 1980, becoming more scattered (Araya-Vergara 1987). This has been attributed to a reduced sandy supply and abnormal heavy rain. Between Rapel and Bio-Bio Rivers ebb and flow deltas are typical (>Fig.  4.3.8). The deltaic estuarine banks become more stable south to the Maule estuary then decrease and become ephemeral down to the Bio-Bio estuary. Tidal deltas are rounded and lobate on the seaward side, where wave energy is higher, and more irregular on the landward side. Off Valparaiso are the Islas Juan Fernandez, where the Scotsman Andrew Selkirk was marooned by his captain, William Dampier, in 1704. Robert Louis Stevenson wrote the story of his survival (as Robinson Crusoe) on these remote and uninhabited Chilean islands, one of which is now called Isla Alejandro Selkirk and another Isle Robinson Crusoe.

Chile

⊡⊡ Fig. 4.3.4 Map of southern Chile.

Copiapó R.

4. 3

First river outlet with permanent runnof after the desert

Elliptic and zeta form bays with progressive surf zone change down drift Tongoy

740 km

72° 31°

Complex intermediate surf zone in front of estuary, development of beaches and dune systems

Choapa R. Los Vilos

Petorca-Ligua Aconcagua R.

Elliptic and zeta form bay

Transition from prograded rias to estuarine deltas

Valparaiso San Antonio

Dated man induced straightening

Maipo R. Sto. Domingo

Dated foredune accretion and cut and fill

⊡⊡ Fig. 4.3.5 Aconcagua estuary, with paired spits.

Zeta-form bays are common in this sector, and show increasing volumes of sand downdrift to the north, where there are three generations of dunes: early Holocene, middle Holocene and presently active (Araya-Vergara 1989). The Gulf of Arauco (>Fig.  4.3.9) is a zeta-form embay-

ment backed by a cliff cut in soft sediment. The width of the beach and surf zones increases downdrift (northward), indicating sand drift towards the Bio-Bio estuary. Similar features are seen in the embayment between Lebu and Quidico. Associated with these are the Arauco dunes, the

215

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Chile

⊡⊡ Fig. 4.3.6 Parallel foredunes built upon beach ridges on the prograded coast at Santo Domingo. The inner dunes were formed before 600 years bp, and a broad swale separates the inner and outer zones.

⊡⊡ Fig. 4.3.7 At Santo Domingo the foredune was cliffed by storm surges in the winter of 1982, and there is a narrow eroding beach.

largest coastal erg in Chile, with small impersistent compound barchans because the climate is humid, and there are annual floods. The Imperial River meanders across a terrace which is inset in an ancient ria. To the south is a ria coast, formed by partial sub­ mergence of river valleys incised into the schistose coast ranges. The intervening sectors are cliffy. The

­ orthernmost limit of direct influence of the Last n Glaciation is marked by a transition from rias to fiords, förde and submarine glacial piedmonts. The floors of fiords have submerged terraces formed by deposition of fine sediment during Late Glacial and Holocene times, and there are steep bordering glaciated slopes (>Fig.  4.3.10). Deltas at the heads of fiords and mouths of side valleys have streams

Chile

4. 3

⊡⊡ Fig. 4.3.8 Estuarine delta.

⊡⊡ Fig. 4.3.9 Transitional zone between estuarine deltas and rias. Zeta-form embayment

Bío-Bío estuary Gulf of Arauco Lebu

Erg of impersistent barchans

Aprox. 480 km

Quidico

Imperial estuary

Valdivia ría and estuary

Cliffs and erosioanal features

41° 75°

RIAS

Bahía Mansa

which discharge clouds of fine sediment that settle on the fiord floor. Sedimentation rates of some decimetres in the last century have been measured. To the south periglaciated coastal slopes descend to the sea, undercut by marine cliffs on exposed sectors. Active periglaciation results in the formation of screes on steep slopes. In the inner coastline of Chiloé Island the key landforms are förde, formed by marine submergence of former subglacial channels. There are many islands of soft glacial drift with coastlines smoothed by erosion, longshore drift and deposition. Tidal currents are strong because of the large tide range. To the south the coastline in front of the Patagonian ice fields is trenched by deep fiords floored by morainic banks, which were formed during the Last Glaciation and Late Pleistocene to Holocene terraces (ponding esplanades) (Araya-Vergara 1999). Above the fiords are the Patagonian ice fields, with calving glaciers at fiord heads. Glacifluvial outwash produces turbidity plumes and deposits fine to coarse sediment in the fiords. The glaciers have been receding. The San Rafael glacier was an ice lobe protruding into the sea in the 1860s, but by the 1990s it had retreated into a fiord. It is notable for its 70-m high calving ice cliffs. The Magellan region was shaped during the last glaciation, producing fiord and piedmont coastlines. The Strait of Magellan originated as a fiord in the western part, which is similar to the fiords in front to the Patagonian ice field, and a set of piedmont lobes in the eastern part with many cliffs cut in soft glacial, but the fiords are smaller.

217

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Chile

⊡⊡ Fig. 4.3.10 Aysen fjord, with a fjord-head delta and turbidity plume. On the left a hanging valley, and side delta, on the right roches moutonnées on the valley side and a pocket beach.

In the eastern part of the Strait of Magellan the tide range is large and intertidal zones, including well-­ developed shore platforms, are wide. Tidal currents are very strong (e.g. 400 cm/s.) in narrow straits, where they contribute to erosion of the coastline. Some valley mouths have small deltas formed in the Holocene.

References Araya-Vergara JF (1985) Sediment supply and morphogenetic response on a high wave energy west coast. Z Geomorphol Suppl. (Bd) 57:67–89 Araya-Vergara JF (1987) Significance of barchans in beach-dune system interactions in Central Chile. Thalassas 4(1):23–27 Araya-Vergara JF (1989) Remnant coastal dunes and their significance in Chilean ergs. Geoško-Plus 1:15

Araya-Vergara JF (1999) Perfiles longitudinales de fiordos de Patagonia Central. Cienc Tecnol Mar 22:3–29 Araya-Vergara JF (2001) Los ergs del desierto marginal de Atacama. Chile Invest Geogr Chile 35:27–66 Marquardt C, Lavenu A, Ortlieb L (2000). Neotectónica costera en el área de Caldera (27°–28°S), Norte de Chile. In: IX Congreso Geológico Chileno, Puerto Varas, Chile Actas 2, pp 582–588 Paskoff R, Luitiño L, Petiot R (1998) Carácter relicto de la Gran Duna de Iquique, Region de Tarapacá, Chile. Rev Geol de Chile 25(2):255–263 Paskoff R, Manriquez H, Cuitiño, L, Petiot R (2000) Caracteristicas, origen y cronologia de los cordones dunares de la playa de Santo Domingo, Region de Valpararaiso. Chile Rev Geol de Chile 37(1):121–131 Potter PE (1994) Modern sands of South America: composition, provenance and global significance. Geol Rundsch 83:212–232

4.4  Argentina

Enrique Schnack . Jorge Pousa . Germán Bértola . Federico Isla

1.  Introduction The coastline of Argentina is about 5,700 km long, with coastal landforms shown in (>Fig. 4.4.1). The coastal geology ranges from a fairly stable area in the north to a generally rising area in Patagonia and tectonically-affected coasts in Tierra del Fuego. A continental shelf, in places, over 800 km wide, is covered by terrigenous sediment accumulated during Quaternary sea level oscillations. Pebbles are a common component of Patagonian beaches, but sandy beaches are typical along the coast from Cabo San Antonio to Mar Chiquita lagoon, and along the southern Buenos Aires barrier coast. Salt marshes are present along the mostly macrotidal Patagonian coast, and on mesotidal areas, with extensive development in Bahía Anegada and Bahía Blanca, where freshwater influence is absent except for river outlets. Brackish marshes with freshwater mixing are found in the northeastern sector of Buenos Aires Province (Samborombón Bay, Mar Chiquita Lagoon). Cliffs are extensive on the Patagonian coast, but the Buenos Aires (Pampas) coastline alternates between low-lying and cliffed areas. Evidence that the sea stood at higher levels dur­­ ing  Pleistocene interglacial phases and the postglacial Holocene is seen on the Argentine coast as beach ridges, marine terraces and estuarine deposits, in most cases hosting a molluscan fauna which has enabled dating of the deposits. In many low-lying areas with marshes evidence of Quaternary emerged shorelines is found, and salt marsh development is strongly related to the postglacial sea level fluctuation. The climate of the Argentine coast is cold and humid in Tierra del Fuego, arid and semi-arid from Rio Gallegos to Bahia Blanca and temperate-humid from there to the Paraná Delta. Predominant wave directions for the Argentine coast are S, SE and E, with wave periods be­­ tween 6 and 16 s, the latter corresponding to ocean swell approaching from the south. There are stormy periods with waves up to 5 m high along the eastern Tierra del Fuego and on the Buenos Aires Atlantic coast. Tides along the Atlantic coastline and on the Rio de la  Plata are predominantly semidiurnal. Tide ranges are

large (megatidal) on the east coast of Tierra del Fuego and along most of the Patagonian coast, with maximum spring amplitudes reaching more than 10 m at San Sebastian Bay (Tierra del Fuego), and about 10 m in Magellan Strait, at Rio Gallegos and at Punta Loyola (Santa Cruz province). Puerto Gallegos has 10.4 m, Puerto Santa Cruz 9.5 m, and Bahia Oso Marino 9.2 m. Mesotidal environments are present from the area north of the Rio Negro outlet (3.3 m at Punta Redonda) to Monte Hermoso, near Bahia Blanca. From there northward the coast is typically microtidal: Mar del Plata has 1.1 m and Buenos Aires 1.0 m. Many storm surges have been recorded along the Argentine coast simultaneously with the northward travelling tidal wave. The duration of these storm surges ranges from a few hours up to 2 or 3 days. They are basically produced by the combined action of an anticyclone located to the west of Argentina (semi-permanent Pacific anticyclone) and a cyclone located over the Atlantic to the SE of Argentina, the latter moving towards the east or northeast. Because of this situation, strong winds from the south or southeast and high water levels affect the whole Argentine coast, as well as the Rio de la Plata shores. Erosion is in progress on the sandy barriers of northern Argentina, particularly when a storm surge coincides with high spring tides. Storm surges are considered the most significant natural agent for coastal erosion on the eastern coast of Buenos Aires, and the Rio de la Plata area. Historical mean sea level trends studied from tide gauge records (Lanfredi et al. 1998) show a rise of 1.6 ± 0.1 mm/ year over 70 years at Buenos Aires, 1.4 ± 0.5 mm/year for Mar del Plata and 1.6 ± 0.2 mm/year for Quequén. Recent calculations for Buenos Aires and Mar del Plata over an extended period show a coincidental rise in sea level of ~1.6 mm/year (Pousa et al. 2007; D’Onofrio et al. 2008).

2.  The Argentine Coastline The southernmost coast of Tierra del Fuego (northern Beagle Channel coast and southern Atlantic coast) extends for 220 km from west to east. It is an indented rocky coastline with pocket gravel beaches in embayments.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4. 4

Argentina

⊡⊡ Fig. 4.4.1 Predominant landforms on the Argentine coast. BUENOS AIRES Río

R

ío Sal ad

de

o

TA

36°s

N

D

IL

VE

IA

NT

BAHIA BLANCA

MAR DEL PLATA

MONTE HERMOSO

0m

gro

Ba hi Saan dAentonio Caleta Valdés

ub Ch

Río

Ne

IA

40°s

d o

Rí o

Colora

Bahía Samborombón Cabo San Antonio

Laguna Mar Chiquita

AN

Río

la Pl ata

20

220

ut

44°s

Deltas Estuaries and marshes

Bahía Bustamante

COMO DORO RIVA DAVIA

Sandy shoreline Pebbly shoreline Cliffs

48°s

Ice-fringed coast

Bahía San Julián Río Santa Cruz

60°s

Islas Malvinas

Río Gallegos

w

74°

25° w

Estrecho de

52° Magallanes

Bahia San Sebastíán

0 100 200 300 km

Isla de los Estados

70°w

66°w

62°w

58°w

54°w

50°w

Argentina

The ­channel is 5 km wide, its average depth between 100 and 450 m. Holocene raised beaches occur at various elevations along the southern coast. The Beagle Channel is a tectonic trough that was completely covered by ice during the Last Glaciation and occupied by a glacial lake at about 9,400 bp. It was invaded by the sea at about 8,200 bp and the marine environment was fully established along the channel by at least 7,900 bp, with a maximum sea level attained between 6,000 and 5,000 years bp (Rabassa et al. 2000; Bujalesky 2000). The Beagle Channel is an estuarine (fiord) system controlled by significant and seasonal freshwater input and by tidal inflow from both west (Pacific) and east (Atlantic). The channel is microtidal with a semi-diurnal regime. Mean tide range is 1.1 m at Ushuaia and the tidal wave moves from west to east. The embayments of the indented rocky coastline of the Beagle Channel are related to tectonic alignments affected by successive glaciations, and are occupied by small gravelly pocket beaches. Along the Beagle Channel Holocene raised beaches stand up to 10 m above mean sea level, and are often capped by anthropogenic shell middens. The oldest Holocene coastal deposits may be partly the result of isostatic recovery while the younger levels are due to recent tectonic uplift. The Bahía Lapataia-Lago Roca valley (20 km west of Ushuaia) is a palaeofjord that was occupied by a lateral and tributary valley glacier during the Last Glacial maximum (about 18,000–20,000 years ago). Well-rounded glacially moulded rocky hills (roches moutonnées), lateral moraines and kames are present in this area and Sphagnum peat bogs occupy the lowlands. Holocene marine depositst are scattered along Bahía Lapataia, Archipiélago Cormoranes, Río Ovando, Río Lapataia and the eastern coastline of Lago Roca, overlying glacial landforms and reaching a maximum altitude of at least 8.4 m above mean sea level (Bujalesky 2000). The eastern (Atlantic) coast of Tierra del Fuego extends for 330 km from SE to NW. It is macrotidal to megatidal, and exposed to high energy waves and strong westerly winds. Extensive wide beaches are composed of gravel and coarse sand. Pleistocene glacial drift deposits form high cliffs in its northern section, and these and other submerged glacial deposits have supplied sediment to beaches. There are gravelly beaches and a narrow Holocene gravelly beach ridge plain (250 m wide) runs between Cabo Domingo and the Río Grande inlet, attached to the base of an Upper Pleistocene marine terrace. In the northern part of the coast San Sebastian Bay is a semicircular embayment (55 km by 40 km) occupying a

4. 4

wide depression formed by glaciers during the Pleistocene and reshaped by the sea during the Holocene transgression. It includes a fossil marsh, the upper marsh, controlled by deflation and dominated by Salicornia sp. scattered on the mudflat, and Spartina grass progressively buried on the windward sides, gravel ridges and cheniers. The mudflat occupies much of the tidal flat. At low tide meandering tidal channels are seen. Península El Páramo is a 20 km long gravel spit barrier that partly shelters San Sebastián Bay. Cliffs extend 40 km from Cabo Nombre (10 m high) to Cabo Espíritu Santo (90 m high) in the northern part of Tierra del Fuego. They are cut in glacial drift deposits and Tertiary silty sands. At Punta Sinaí is an extensive erratic boulder field, and erratics can be seen on the beaches. In Patagonia most of the estuaries (with the exception of the Chubut, Negro and Colorado rivers) have low discharges, and tidal processes are very active at their outlets. Beaches are typically gravel-dominated, showing at times several ridges on the backshore and on the beach face (>Fig. 4.4.2). In some protected embayments, the coarse substrate carries marsh vegetation (>Fig. 4.4.3). Patagonian gravels extend over the whole region. They have been reworked by marine action at various higher sea levels and are typical of the Pleistocene and Holocene raised beaches as well as the modern shore deposits, which contain varying proportions of sand. Much of the coast has cliffs (>Fig. 4.4.4), cut in Tertiary continental and marine deposits and Jurassic volcanic rocks, and there are many embayments. Evidence of Pleistocene and Holocene high sea level stands extends along the coast, where six marine terraces, at least three of which are of Quaternary age, have been recognised. The most recent (Holocene) reaches an altitude of 8–12 m above mean sea level and extends from eastern Tierra del Fuego to northern Patagonia. The coast of Península Valdés has cliffs and shore platforms cut in Miocene sedimentary rocks, and the Caleta Valdez lagoon is bordered by a series of Pleistocene emerged gravelly beaches (Schellmannn and Radtke 2003). Marshes are found in some protected low energy environments. Except for the northern (Río Negro-Bahia San Blas) marshes, which are mesotidal, the Patagonian wetlands are mostly macrotidal. They are present either in estuaries (Chubut, Deseado, Santa Cruz, and Gallego Rivers, and in NE coast Tierra del Fuego) or within sheltered coastal bays. Some of the marshes are eroding, exposing peaty microcliffs. San Antonio Bay, which has a maximum spring tide range of 9 m, is bordered by pediments, hollows, Pleistocene and Holocene coastal and marine deposits and

221

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4. 4

Argentina

⊡⊡ Fig. 4.4.2 Gravel beaches and berms near Bahía San Julián (Santa Cruz, Patagonia).

⊡⊡ Fig. 4.4.3 Salt marsh vegetation on a gravelly beach near Bahía San Julián.

marshy environments (Schnack 1985). Extensive sandy shoals are exposed at low tide at the outer part of the bay. There are slumping cliffs up to 8 m high in Miocene sediment at Las Grutas.

The coast from San Antonio Bay to the mouth of the Rio Negro has alternations of beach and dune-fringed low coasts and cliffs. The Rio Negro estuary has a mesotidal regime. Two shoals (Banco Miguel and Banco La Hoya)

Argentina

4. 4

⊡⊡ Fig. 4.4.4 Cliff and abrasion platform at Peninsula Valdés.

form an ebb delta composed of well-sorted fine sand. The assymetry of the delta is mainly due to northward longshore drifting. From the Rio Negro estuary to the Rio Colorado delta the coast is predominantly low-lying, except for cliffs immediately south of San Blas. Pebbly beach ridges of Holocene and Pleistocene age are also present near San Blas. The Rio Colorado is generally regarded as the northern limit of coastal Patagonia. The south Buenos Aires (Pampas) coast extends from Bahia Blanca to the Paraná river delta. The mesotidal Bahía Blanca estuary is formed by several channels oriented NW-SE and separated by extensive tidal flats and islands. The area between Bahía Blanca and Bahía San Blas may have formerly been part of the delta of the Río Colorado, and the tidal flats are derived sediment. Mesotidal beaches, typically sandy, extend from Monte Hermoso towards Miramar, alternating with cliffs (Isla et al. 1996). In low-lying areas, such as Claromecó and Quequén there is evidence of Pleistocene and Holocene high sea levels. North and south of Mar del Plata high cliffs cut in Plio-Pleistocene sediment are retreating rapidly, at times reaching 1–2 m/year. At Mar del Plata the cliffs are in more resistant Lower Palaeozoic quartzites. From Mar Chiquita beach and tidal inlet north to Punta Rasa is a low-lying coastal plain, with dunes on a major coastal barrier. The geological substratum of this segment consists of Holocene estuarine and marine deposits, although evidence of former Pleistocene sea levels is also found.

At Mar Chiquita these deposits are well developed, with shelly beach ridges and the lagoon marginal flat built on the estuarine deposits. Following a maximum sea level stand of about 2.5 m above present level about 6,000 years ago the sea fell, and a southward-prograding barrier enclosed the Mar Chiquita lagoon. The tidal inlet has shown northward migration interrupted several times by human intervention. Severe erosion is occurring at Mar Chiquita beach immediately south of the inlet. Extensive dunes along the coast north of Mar Chiquita are in many places fixed and urbanised. Wave energy generally diminishes along the coast from Mar del Plata to San Clemente del Tuyú, but at Punta Médanos larger waves are due to the nearshore presence of linear sand ridges oriented SSE 30° to the coastline, which concentrate wave energy. These ridges continue south beyond Pinamar. At Mar de Ajó mean wave heights are 0.68 m, whilst maximum annual wave heights are 1.31 m. Erosion has lately dominated the Buenos Aires coastline northward to the Mar Chiquita lagoon. The southern beaches (Mar del Plata, Santa Clara, Mar de Cobo) are typically reflective, while the beaches at Villa Gesell, Pinamar and Partido de la Costa generally have a dissipative profile with the development of sand bars. In those areas where there is no human pressure, as in Mar Azul (south of Villa Gesell), or where sediment transport is blocked by a hydraulic barrier, as at San Clemente del Tuyú, beaches are accreting with backshore dunes (>Fig. 4.4.5). At Mar del Plata and other sites, the emplacement of sea walls, jetties and piers is responsible for erosion along the littoral drift direction.

223

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Argentina

Dune fixation by vegetation, traditionally considered beneficial, has been claimed to be responsible for alteration of the sediment supply to the beach. A transitional zone extends for more than 100 km at Samborombón Bay, in the central part of the Salado depression and part of the Río de la Plata system. The

hinterland is a low-lying portion of the Pampas Plain, occupied by Pleistocene and Holocene coastal and marine deposits. At Samborombón Bay a brackish marsh is present behind an extensive mudflat. From Punta Piedras to Buenos Aires the Río de la Plata coast runs SE-NW and the coastal plain is up to 8 km wide. rising to about 5 m. ⊡⊡ Fig. 4.4.5 Stable or accreting beach on the NE barrier of Buenos Aires at Mar Azul, south of Villa Gesell.

⊡⊡ Fig. 4.4.6 Storm surge damage at Mar del Tuyú.

Argentina

It  formed during the mid-­Holocene maximum sea level stand around 6,000 years bp (Cavallotto et al. 2005) and the subsequent emergence. The coastal plain is composed of Holocene beach ridges, marshes and fresh-water wetlands, with a levee of estuarine sediment on the coastal fringe. It is backed by a high terrace on continental sediment, mainly loess-like silts. The inner boundary of the coastal plain is the midHolocene high sea level coastline. The Río de la Plata has an area of about 35,000 sq km and a drainage basin (Paraná and Uruguay Rivers) of more than 3 million sq km. The estuary receives large amounts of dissolved and suspended solids and organic carbon. The sediment is largely from the Paraná River, mainly as suspended load but also as bed load, and forms a delta of 18,000 sq km. The sediment in the estuary consist mainly of silty clays with some sand from the Paraná river, while the yield From the Uruguay River is sandy. Coastal plain flooding due to storm surges can be particularly destructive in areas where topographic gradients are extremely low, as on the Río de La Plata shores and in the Salado basin. Storm surges are regarded as the most important morphodynamic factor in coastal development along this area (>Fig. 4.4.6). It is widely accepted that a global sea level rise is causing impacts on coasts, and an accelerated sea level rise would certainly exacerbate coastal erosion and flooding in Argentina. The most vulnerable areas are the Río de la Plata, including the Paraná delta and the estuarine coastal plain, already exposed to frequent storm surges, and the sandy coastline extending from Mar del Plata to Punta

4. 4

Rasa, close to Samborombón Bay, where there is already severe erosion. It is likely that other areas such as Samborombón Bay and some estuaries would be affected by an increasing sea level, depending on its magnitude and the interaction with other, human and climatic variables.

References Bujalesky G (2000) Quaternary Coastal Environments of Tierra del Fuego (Argentina). Field Trip Guidebook, November 4–7. IGCP-437, IGU, INQUA, p 27 Cavallotto JL, Violante RA, Parker G (2005) Sea level fluctuations during the last 8,600 years in the La Plata river. Argentina. Quat Int 114:155–165 D’Onofrio EE, Fiore MME, Pousa JL (2008) Changes in the regime of storm surges at Buenos Aires. Argentina. J Coastal Res 24:260–265 Isla FI, Cortizo L, Schnack EJ (1996) Pleistocene and Holocene beaches and estuaries along the Southern barrier of Buenos Aires. Argentina. Quat Sci Rev 15:833–841 Lanfredi NW, Pousa JL, D’Onofrio EE (1998) Sea-level rise and related potential hazards on the Argentine coast. J Coastal Res 14(1): 47–60 Pousa JL, Kruse E, Guaraglia D, Mazzoldi A, Carbognin L, Tosi L et al (2007) Geological hazards in two sandy environments: the eastern coast of Buenos Aires (Argentina) and Vence (Italy). Env Geol 51:1307–1316 Rabassa J, Coronato A, Bujalesky G, Salemme M, Roig C, Meglioli A et al (2000) Quaternary of Tierra del Fuego, Southernmost South America: an updated review. Quat Int 68–71:217–240 Schellmann G, Radtke U (2003) Coastal terraces and Holocene sea level  changes along the Patagonian Atlantic coast. J Coastal Res 19:983–996 Schnack EJ (1985) Argentina. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 69–78

225

4.5 Uruguay

1.  Introduction The coastline of Uruguay is about 600 km long, including the shores of Rio de La Plata and the Atlantic Ocean. Evolution of the coastline has been influenced by the uplift of the Ballena and Animas ranges and the Southeast and Southwest highlands, and subsidence of the structural trough that underlies the Santa Lucia River valley and Merin Lagoon basin, and lowlands occupied by Montevi­ deo Bay and East Coast lagoons and estuaries. In the eastern part of the country, the coastline truncates the geological formations of the Eastern Highlands Shield, which shows a north–south trend (N 10° E) in the folded and faulted rocks and the pattern of granite intrusions. The Pre-Devonian crystalline rocks of the Brazilian craton to the west extend to the south, in the Ingles and Rouen banks (Rio de La Plata), where they dip north into the Rio Salado syncline (Bossi 1966). North of the coastal region, the north–south trend of the geological formations of the Eastern Highlands is interrupted by a large complex of strike-slip faults, with a lateral displacement of 30 km, and a general direction of N 60° E between the tectonic basin of Merin Lagoon and the Arazati River mouth. These strike faults have led to vertical uplift of local areas between alternating areas of tension and compression. This fault zone has been filled in succession with pillow lava, rhyolitic dykes (Tertiary volcanism), and sea-floor, coastal, and continental sediment. The coast is considered to have been generally stable in late Quaternary times, indicating a higher sea level during the Holocene. Earthquakes, however, are occasionally recorded and the possibility of localized minor warping cannot be ruled out. The resultant general direction of the shoreline is northwest–southeast between Gorda Point and Colonia Point and WNW–ESE between Colonia Point and the Santa Lucia River estuary. East of Espinillo Point to Del Este Point, there is an irregular eroded rocky coast, with many bays and beaches, and a long coastal arc tending generally in a west–east direction. East of Del Este Point to Santa Maria Cape, the

general coastal direction is N 60° E, with a gradual turn to N 30° E in three sectors of the coast up to the Brazilian frontier. The variations in coastal aspect are important for coastal dynamics. East of Espinillo Point the coast is exposed to a SE ocean swell, modified by the shallow depths of Rio de La Plata and the epicontinental sea, and to storm waves mainly from the SE and SW (mouth of the Rio de La Plata). The western sector is subject only to storm waves, mainly from the SE and SW, occurring in a reduced fetch and shallow water (5 m maximum depth). Mean spring-tide ranges on the coastline are generally between 0.4 and 0.6 m, but the SE winds produce storm surges between 1.9  m (Rio de La Plata mouth), 2.7  m (Montevideo), and 3.7 m (Colonia). Cliffs are found on several geological formations and in many sections are fronted by beaches. Sandy beaches are most common, fronted by sand bars and backed by beach ridges and dunes. The sand is predominantly quartzose, but calcareous in the east and the far west. Longshore drifting on the ocean coast north of Santa Maria Cape is generally from SW to NE. There is no net drift between Santa Maria Cape and Del Este Point, but west of Del Este Point there is some westward drifting. The shore between Punta Ballena and Punta Rubia has the coarsest beach sand (1 mm); elsewhere it is finer (0.1–0.2 mm) (Jackson 1979). Most Uruguayan rivers have occasional torrential flow and drain along fault-guided valleys into estuarine lagoons, with sea entrances blocked or encumbered by inwashed sandy thresholds and flanking spits. Swampy shores, typically with Juncus and Spartina salt marshes, occupy the more sheltered parts of the coast between Montevideo and Arazati, the margins of estuaries, and the mouths of coastal lagoons.

2.  The Uruguayan Coastline The Argentine border runs N–S down the Rio Uruguay to the Rio de la Plata gulf. The Rio Uruguay widens below

Edited version of a chapter by J. M. Jackson in The World’s Coastline (1985: 79–84). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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4. 5

Uruguay

Fray Bentos to an estuary with numerous alluvial islands. Below Fray Bentos it is about 6 km wide, with the delta of the Rio Negro on its eastern shore, but it then narrows as it is constricted by the delta of the Parana River, coming in from the west. It widens again past Carmelo, the eastern shore passing from estuarine to marine and sandy along the Colonia coast. This coastal lowland is fringed by beaches of coarse sand brought down by the Uruguay River. Gently curved beaches between cliffed promontories of Miocene–Oligocene mudstone are backed by bluffs cut in Pleistocene sandstones and covered with dunes. The coast curves out to a granitic promontory, Punta Hornos, with outlying islands and rocks, and the port of Colonia del Sacramento. Tertiary rocks outcrop above the granite in low cliffs on Colonia Point, continuing as the coastline turns eastward, and the Rio de la Plata gulf diverges as an area of shallow fresh water some 40 km wide. The mouth of Rosario River is navigable, but there are broad sand bars offshore. To the east the beach is interrupted by headlands and shore platforms of granites with diabase dykes, then a broad lowland with fresh-water swa­ mps and dunes has developed here in front of a Pleistocene coastline, which is indicated by bluffs. There are low-lying plains with dunes and bordering cliffs in Tertiary formations near Arazati offshore sand shoals which are a source of continuing beach progradation, in contrast to the prevalent erosion in the San Gregorio sector to the SE. There is westward drifting of beach and shoal sands towards Jesus Maria Point, and the sandy beaches are backed by salt marshes with Spartina. Near the port of Juan Lacaze is the small Rio Rosario delta, and a cliffed coast runs out to Punta San Gregorio and on to Punta del Tigre. The 40 m high cliffs on Punta San Gregorio has been cut into Plio-Pleistocene limestones and sandstones, and sand derived from these has formed sand shoals that are the source of continuing beach progradation to the east, where 30 km of sandy barriers and spits formed in Pleistocene times fill an ancient bay. The Rio Santa Lucia flows by way of a small delta into a bay where beach ridge plains fringe the coast. Structurally controlled by faults, the Santa Lucia River valley borders a scarp on its eastern margin. The Santa Lucia Delta plain is developed on a swamp area, formerly a chain of coastal lagoons, that occupies a marine embayment at the southern end of the Santa Lucia tectonic depression. On the eastern side of Punta del Espinilla a succession of bays between small headlands extends to Montevideo Bay, which occupies a sunkland between the San Pedro Peninsula to the east and the Lobos Peninsula to the west.

The port of Montevideo is sheltered by large breakwaters. In the NE of the bay some inter-tidal mudflats are threatened by erosion as a consequence of dredging a navigation channel, and sandy beaches on the eastern shore (as at Capurro) were destroyed by port construction. The sand deposited in these bays came from the sea floor during and after the Late Quaternary marine transgression, but there is no longer any significant sand supply from this source. The coast to the east consists of low cliffs and coves with beaches, the largest of which is Playa Ramirez. There are steep slopes rising to Montevideo Hill, and small bays with sandy beaches are backed by partly mobile dunes. Punta Manso is the easternmost of the amphibolite and pegmatite headlands characterising the rocky coast of Montevideo. Brava Point also consists of Pre-Devonian rocks and amphibolite and pegmatite dykes dissected by marine erosion into irregular shore topography, and embayments fringed by Spartina and Juncus marshes. On either side of Brava Point, valley mouths end in sandy coves backed by dunes. East of Montevideo the southerly fetch over the South Atlantic Ocean lengthens, so that wave energy increases. Seaside resorts continue to Atlantida and round the bay to Piriapolis. There is a large beach of fine Pliocene and PlioPleistocene rocks backed by mobile quartzose sand dunes (PCYMP 1979) between Punta Manso and Punta Negro. The Animas Range, of Pre-Cambrian igneous rocks, runs south to the coast to form the promontory out to Punta Negra. The coastline has a jigsaw configuration resulting from the diagonal trend of the faults (strike-slip and oblique slip faulting) leading to movement of the coastal blocks. To the east are landslips in Tertiary formations that rise toward the Animas Range. This formation sometimes has steep coastal slopes with only limited basal cuffing behind shore platforms. The shore platforms are cut in hard Paleogene sandstones containing pebbles and cobbles, which are was­ ­hed out to form beaches in the coves. Behind the narrow Solis Grange River entrance, impeded by looped sand-and-pebble bars and a spit that has grown from west to east, there is an estuarine lagoon surrounded by salt marshes. East of this mouth the beach is backed by limestones and peaty deposits capped with dunes. The peats were formed in an estuarine ­environment. Seaward, these Holocene sediment have been cut back into steep cliffs (4–6 m) and serrated shore platforms. The same pattern characterises all the river mouths in this coastal sector. Punta Negra is formed by three syenitic ridges with diabase dykes, bordered seaward by extensive rocky

Uruguay

­ utcrops. The coves between the three ridges have beaches o of coarse sand. East of Punta Negra is an asymmetrical sandy bay with a beach of fine sand backed by dunes, which have drifted northeastward to the Laguna del Sauce, a fresh water lake that occupies a tectonic depression and is linked to the sea by a narrow channel. Ballena Point is a promontory of quartz sandstone, dipping 80 west and ending in a rocky slope that plunges into deep water, and another asymmetrical bay curves out to Punta del Este, south of Maldonado. It has a permanent natural outlet opening over a granite substratum, but this is not navigable due to a perennially recurring looped sand bar offshore. Offshore is a high island, Isla de Lobos, with a lighthouse. Punta del Este is a breached sandy tombolo in the lee of a granitic island. Settlers built a road connecting the island with the mainland around 1,900, making this a permanent isthmus. Bluffs back the tombolo, which is covered by mobile dunes. Beyond Punta del Este the coastline trends ENE following the trend of schist outcrops, and sandy beaches receive strong southeasterly ocean swell. The beach has a 10° slope and consists of coarse quartzose sand derived from weathered granite. Maldonado Inlet is a broad, shallow estuarine lagoon with extensive sandy shoals exposed at low tide, and shores that are fringed by salt marshes. There has been extensive sand deposition, with dunes surmounting barrier formations that contain Laguna José Ignacio, Laguna Garzon and Laguna de Rocha, which has a swampy deltaic northern shore. The three lagoons occupy former marine embayments in tectonic depressions, enclosed by sand barriers in Pleistocene times. The barriers were dissected during the last glacial low-sealevel phase, before the Holocene outer barrier was added to en­­close the present lagoons. The lagoons have intermittent natural outlets. There is a granite promontory and some minor granite headlands that interrupt a 75-km-long beach backed by bluffs cut into Pleistocene coarse-grained calcarenite. The barrier bears dune ridges parallel to the coastline, and parabolic dunes and blowouts advancing SW–NE. Cabo Santa Maria is a complex tombolo formed in the lee of schist outcrops and is covered by mobile dunes, which have spilled northeastward. Bluffs back the tombolo. The coastline then trends NE past La Paloma in a series of long asymmetrical bays, shaped by refracted ocean swell, with beaches backed by extensive dune ridges. The beach is in­­ ter­­rupted by outcrops of chloritic–schist phyllite running N 25° E and dipping 80° north, which form the Punta Rubia promontory.

4. 5

Rivers converge into Laguna de Castillos, which has a 22 km long meandering outlet channel through the dunes to the coast north of Punta Aguda, where granite outcrops on the shore. The lagoon is bordered by emerged Holocene beaches (2–4 m above sea level) and fresh-water swamps. Brackish water flows in through the channel from time to time. South of this beach the granitic headlands of Polonio are covered by mudstone and peat, overlain by fixed dunes composed of coarse quartzose sand capped by a recent longitudinal mobile fine dune. The deposits are up to 60 m thick above the granite. In places dune sand is spilling down on to the shore. Recent dunes also found around swales between the Pleistocene calcarenite ridges which back a 25 km long, gently curved beach. These form a barrier impounding Laguna Negra, the dunes continuing behind low cliffs and rocky shores on the granite outcrop at Punta Palmar behind the Coronilla islands. This is the Santa Teresa Granite, part of an intrusive ring pattern. Granitic headlands separate gently curved arcuate beaches of quartzose sand backed by dunes and bluffs of Pleistocene dune calcarenite. In the lee of the Coronilla Islands, wave refraction has shaped a major sandy breached tombolo. The SE Santa Teresa coast has a long sandy beach with scattered gravel. The long dune ­calcarenite ridge has been much dissected by lavakas (­gullies). Runoff playas are transformed into mobile long­­ itudinal dunes of fine sand by northerly and southwesterly winds. The gently curving sandy Chuy Beach is backed by extensive, partly mobile, quartzose-calcareous dunes, and two long Pleistocene calcarenite ridges cut by gullies and lavakas. Between the ridges are fresh-water lakes and associated swamps, underlain by rhyolite dykes and Santa Teresa granite. The ridges form a barrier on the southern side of the Merin-Patos Lagoon, which extends into Rio Grande do Sul, and the beach runs along the Brazilian border at Barra de Chuy.

References Bossi J (1966) Geologia del Uruguay, Departamento de publicaciones de la Universidad, Uruguay Jackson J (1979) Sedimentacion reciente en la costa platense-atlantica del Uruguay al este de Montevideo, Informe Ministerio de Obras Pitblicas J Uruguay

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4.6  Brazil

Dieter Muehe

1.  Introduction The Brazilian coastline (> Fig.  4.6.1) lies between latitudes 4°N and 32°S and is about 7,400 km long, although various estimates range between 6,000 km and more than 10,000 km (Muehe 2003). It is backed by extensive continental plateaux of igneous and metamorphic PreCambrian rocks with basins of Palaeozoic sediment, and fronted by a continental shelf that widens in the North, in front of the Amazon. Sedimentary terraces, pediments of Tertiary age, of varied composition (Barreiras Group), extend from the North to the Southeast regions and represent an important element of the coastal landscape. The flat topped surface was deeply incised during Pleistocene low sea levels while the distal portions of the deposits were eroded by the sea during marine transgressions, leaving a coastline of fossil and active bluffs several metres high, preceded by narrow beaches. Rocky sectors are found mainly in the Southeast region, in places where the Pre-Cambrian basement structures of the Serra do Mar mountain range come to the coastline, as in the area around Ilha Grande bay (Rio de Janeiro) up to Caraguatatuba and São Sebastião island (São Paulo). Mangroves grow alongside the estuaries and are highly developed in the North region, between French Guiana and the Maranhense Gulf, comprising the States of Amapá, Pará and particularly Maranhão, which has about half of the area of Brazil’s mangroves. Large dune fields are found in the Northeast region where the climate is dry, and also in the Southeast and South regions in places with strong winds and a supply of fine sand. There is wide development of dunes along the coast of Rio Grande do Sul, in front of the Patos lagoon, as also locally in Espirito Santo, Rio de Janeiro and Santa Catarina States. > Fig.  4.6.1 is a general geomor­ phological chara­cterisation based on classifications of Silveira (1964); Dominiguez (2004) and Muehe (2006).

The climate varies from warm temperate in the south, where the winters are cool and rainy, to humid tropical in the north. Porto Alegre, in the south, has a mean monthly temperatures ranging from 14.3°C to 24.1°C, with precipitation well distributed through the year, amounting annually to about 1,400 mm. In Rio de Janeiro in the southeast the mean temperature of the coldest month is 20.6°C in July, rising to 25.6°C in January, and there is an average annual rainfall of 1,082 mm with a relatively dry winter. Recife, in the northeast, has a mean annual temperature of 25.4°C with monthly mean temperatures varying less than 2°C around this mean. Annual mean precipitation is slightly less than 2,500 mm with major rainfall between April and July. Belém in the north has a mean annual temperature near 26°C with little variation through the year. Annual rainfall is 2,820 mm with a wet season from January to May, when mean monthly rainfall is above 300 mm and a less moist period during the rest of the year with monthly rainfall between 100 mm and 200 mm. The west-facing coastline (north west of Cabo Cal­ canhar) receives waves generated by east and southeast trade winds, but wave energy is low in the sector of the Amazon and Pará estuaries, where the continental shelf is wide and shallow. Farther south, waves generated by the trade winds are effective, and they are accompanied by southerly swell arriving from the South Atlantic. Wave energy increases in the stormier environments of the far south. Tide ranges are very small (less than 2 m) in the south, southeast and part of the northeast, up to the state of Alagoas. Further north tide ranges of up to 12 m and tidal currents of up to 250 cm/s have been recorded in the Amazon estuary. From south to north mean spring tide ranges are: Rio Grande 0.3 m, Imbituba 0.7 m, Santos 1.1 m, Rio de Janeiro 1.1 m, Cabo Frio 0.9 m, Salvador 2.2 m, Maceió 2.1 m, Recife 2.1 m, Natal 2.3 m, Mucuripe 2.5 m, São Luis 5.2 m, Belém 2.2 m.

This is an abridged version of a chapter by O. Cruz; P.N. Coutinho; G.M. Duarte; A. Gomes and D. Muehe which was published in The World’s Coastline (edited by E.C.F. Brid and M.L. Schwartz) in 1985 pp. 85–91.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 4.6.1 The coastline of Brazil.

2.  The Brazilian Coastline From the Uruguay border the coast runs northeast (> Fig. 4.6.1A) and consists of wide sandy barrier formations enclosing major lagoon systems, Lagoa Mirim and Lagoa dos Patos, with associated salt marshes and sedge swamps (Delaney 1965). The coastal plain of Rio Grande do Sul is up to 120 km wide, and notable for extensive dunes driven by predominantly northeasterly winds. The coastal barrier fronting the Lagoa dos Patos has dunes up to 25 m high in a 10 km wide zone along its seaward fringe, disrupted by blowouts that have moved from NE to SW. This is backed by a dune ridge that rises to 35 m, and has been dissected, yielding sand for the cuspate forelands between excavated bays on the shores of Lagoa dos Patos. The sandy ocean beach extends almost continuously for 750 km from the southern border up to cape Santa Marta. North from the Patos lagoon to Cape Santa Marta, the scarped edges of the sedimentary and basaltic Paraná basin develop behind a narrow coastal plain with a sequence of lagoons dammed by the barrier beach.

North from Torres to Santa Catarina Island the PreCambrian basement comes to the coast with high promontories separated by valley mouth inlets formed by Holocene marine submergence. The hilly island of Santa Catarina is noted for its beach ridge and dune systems. The southernmost limit of mangroves in Brazil is at Lagu­­­na (28°30'S), with Avicennia ­schaueriana, Laguncula ria racemosa and Rhizophora mangle (Schaeffer-Novelli, Personal communication). North from Santa Catarina island the coast becomes simpler in outline, interrupted by inlets at the mouths of the steep valleys on the eastern flank of the Serra do Mar or by large bays such as Para­ naguá. Pleistocene and Holocene barriers enclose narrow lagoons and mangrove swamps at Cananéia, north of the hilly Ilha do Cardoso (Bigarella 1978) (> Fig. 4.6.2). Stages in the evolution of these depositional features have been traced and dated by Suguio and Martin (1978). The tide range is less than 1.5 m, but behind Ilha Comprida intertidal mangroves are extensive in the lagoonal channel system of Cananéia-Iguape, including the Ribeira de Iguape river mouth. Sandy beaches towards Santos are interrupted by steep-sided residual coastal massifs.

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⊡⊡ Fig. 4.6.1A The coastline of southern Brazil. (Courtesy Geostudies.)

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Santos stands on a peninsula between broad tidal inlets fringed by extensive marshes and mangrove swamps. A curving sandy beach faces southward to the Baia de Santos, between high, steep-sided promontories. At the western end of the beach is a high island attached to the mainland by a sandy tombolo. Sandy beaches are interrupted by rocky promontories that become larger and more frequent as the front of the mountain range of Serra do Mar approaches the coastline. A mountainous promontory runs out toward the high, steep-sided Ilha de São Sebastião, which stands beyond a deep, narrow strait, and the high hinterland dominates the coast between here and Rio de Janeiro. Embayments, some well sheltered from prevailing southeasterly wave action, have beaches and mangrove swamps. Caraguatatuba Bay faces east and is backed by a Pleistocene and Holocene beach ridge plain. The beach prograded following the 1967 downpour in the high hinterland, when catastrophic erosion caused rivers to discharge vast quantities of sediment into this bay (Cruz 1974). Smaller bays with beach ridges are found between steep irregular headlands on the coast as far as Ubatuba, and in the western part of the Baia da Ilha Grande

s­ heltered bays are backed by mangrove swamps, as at the historic town of Parati on the Rio Perequeaçu delta. The indented steep coastline of this bay has rocky headlands (> Fig. 4.6.3) fringed by numerous islands. There are steep, arcuate beaches in small coves, and longer, flatter less regular beaches in protected bays, as in the eastern part of Bahia da Ilha Grande. East of Baia de Ilha Grande the first of a series of large beach barriers, the Restinga da Marambaia, encloses Sepetiba bay. These barriers are in the form of long and straight sandy deposits up to Cabo Frio, interrupted by steep promontories of crystalline rock (> Fig. 4.6.4). Copacabana Beach (> Fig.  4.6.5) stands in front of a heavily urbanized relatively narrow coastal plain, and was artificially nourished in the early 1960s (Muehe and Neves 1995). The Sugar Loaf stands at the western entrance of Baia de Guanabara. East of Rio de Janeiro, the irregular coastline of the coastal range has been rectified by long and straight beach barriers which enclose a series of coastal lagoons, the largest of which is the Lagoa de Araruama, close to Cabo Frio. This is backed by sedimentary formations of the Barreiras Group and crystalline

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⊡⊡ Fig. 4.6.1B The coastline of Northern Brazil. (Courtesy Geostudies.)

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⊡⊡ Fig. 4.6.2 The mountainous coast south of Cananéia. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.3 Rocky headland near Bertioga, showing intertidal oyster zone. (Courtesy Geostudies.)

rocks. There is generally an inner barrier, related to the latest Pleistocene transgression, and an outer barrier developed during the Holocene sea level rise. Muehe and Corrêa (1989) have shown that the barriers migrated to his present position in association with a sea level rise. Lagoa de Araruama is segmented by large spits formed of sediment delivered from the erosive recession of the back of the Pleistocene barrier beach (> Fig. 4.6.6).

Cabo Frio is a dry and relatively cool region, due to the upwelling of cold water by the strong northeastern trade winds. The high evaporation is used for extensive salt production from pans excavated in marshlands or impounded in shallow lagoon areas. The hilly Cabo Frio peninsula is attached to the mainland by a broad sandy isthmus bordered by curving beaches and surmounted by large dunes (> Fig. 4.6.7).

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⊡⊡ Fig. 4.6.4 Rocky promontory at Saquarema, with 1630 cathedral. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.5 Copacabana Beach, Rio de Janeiro. (Courtesy Geostudies.)

North from Cabo Frio the first major feature is the deltaic protrusion at the mouth of the Paraiba do Sul river, formed largely by Holocene beach ridge progradation. The southern part, near Macaé, truncates widely spaced, subdued Pleistocene beach ridges. The large mountain range of Serra do Mar recedes inland and is breached by several river valleys that open on to the depositional coastal plain. Locally, outlying hills reach the coast alongside deep estuaries, as at Vitória (Espirito Santo). Farther

north the large coastal plain of the Doce river is formed of Pleistocene and Holocene terraces and a fringe of beach ridges. The coastal plain widens northward into Bahia (> Fig. 4.6.1B), and includes deltaic regions at the mouths of several rivers, notably the Jequitinhonha, which has a delta at Belmonte flanked by prograded beach ridge plains. A wide beach ridge system marks stages in the growth of the cuspate foreland at Ponta da Baleia in the lee of the

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4.6

⊡⊡ Fig. 4.6.6 The Punta do Acaira spit, projecting into Lagoa de Araruama. (Courtesy Geostudies.)

⊡⊡ Fig. 4.6.7 Blowout in the dunes on Cabo Frio. (Courtesy Geostudies.)

Abrolhos archipelago. The coastal plain then narrows past the mouth of the Rio Alcobaça. Long straight beaches and narrow barriers extend north to Ilhéus on a small delta at the mouth of the Rio Cachoeira. Farther north, transverse faulting has complicated coastal topography north of Itacaré, where the coastline becomes indented by estuaries and embayments, the largest of which is the Todos os Santos Bay, behind the hilly peninsula of Salvador. In an increasingly warm and wet climate, the sheltered shores of the estuaries are extensively mangrove fringed, while the

consistent trade winds produce the waves in coastal waters that have built up beach ridge plains. North from Salvador the relatively straight coast line is bordered offshore by beach rock and calcareous reef sandstones (Bigarella 1975), some of which are cemented dune sands (aeolian calcarenites). There are coastal terraces of 2–3 m and 7–8 m above present sea level. Locally there are cliffs cut into Pliocene sedimentary formations of the Barreiras Group. North of Aracajú the coastline curves out to beach ridge fringed plains around the mouth of the

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São Francisco river; again, there are relics of a Pleistocene delta fronted by prograded Holocene beach ridges. Humid tropical conditions, with annual rainfall between 1,000 and 2,000 mm, dominate the coastline northward, up to Recife and Natal, but with a lengthening summer dry season. The coastline remains simple, with beach ridge plains and nearshore sandstone reefs. Cape Calcanhar marks a major change in the coastline toward a north-­ facing equatorial sector. The coastline is generally smooth and depositional, beach ridges and dunes alternating with lagoons, swampy sectors and salt deposits. The dry season becomes pronounced, and vegetation is sparse in the hinterland. Near the Parnaiba delta are long sandy beaches backed by dunes. Strong onshore winds, dry climate and fine sand have formed the wide dune fields along the coast of Maranhão, west from the Maranhense gulf to Cape Calcanhar, including the whole coast of Ceará and part of the state of Rio Grande do Norte. The Lençóis Maranhenses, a wide dune field located between the Maranhense gulf and the Parnaíba delta, extends along a 20 km coastline up to 10 km inland. Large dune fields also are found along the coast of Ceará as for instance in Jericoacoara where exceptionally huge dunes are developed. Tide range increases westward along this coastline, and tidal currents are strong in the bays of Marcos and São José (Maranhense Gulf) at São Luis, the capital of Maranhão. The coastline in this equatorial sector is indented, with many inlets and estuaries bordered by mangrove swamps and backed by a low plateau of the Tertiary Barreiras Group and early Quaternary sedimentary rocks, which reach the coast in several localities as small cliffs. West from the Maranhense Gulf, the branching funnel-shaped estuary of the Marajó bay near Belém, has strong tides with intricate and variable shoal and channel topography. Numerous mangrove-fringed alluvial islands, and swampy shores are subject to rapid changes in configuration due to interactions of waves and of tidal and fluvial currents. It is separated from the similar Amazon estuary by the large island of Marajó, notable for its extensive swamps and intersecting tidal channel systems. The vast discharge of water and sediment (sand, silt and clay) from the Amazon results in

rapid accretion in and around the mouth of the river and along the adjacent coastline, especially northward into French Guiana as the result of tidal movements and the Guiana current. The low lying coast of the Amazon deltaic region consists of extensive swamps with mangroves and tropical forest, sandy beaches and cheniers, alluvial grasslands, lagoons and numerous channels. The climate is perennially hot and wet, with an annual rainfall of more than 2,000 mm, and vegetation is luxuriant. Coastal waters are turbid due to the high sediment suspension. The low lying swampy coastline continues north to Cabo Orange at the border between Brazil and French Guiana.

References Bigarella JJ (1975) Reef sandstones from Northeastern Brazil. A Ac Br Ciên 47:395–409 Bigarella JJ (1978) The Serra do Mar mountain range in the Eastern part of the State of Paraná, SEPL/ADEA, Curitiba, 248p Cruz O (1974) The Serra do Mar mountain range and the coast of Caraguatatuba (in Portuguese). Dissertation, Instituto de Geografia – University of Sao Paulo, 11. Sao Paulo Delaney PJV (1965) Physiography and surface geology of the coastal plain of Rio Grande do Sul (in Portuguese). Escola de Geologia. UFRGS, Special Edition, Porto Alegre 6, 105p Dominiguez JML (2004) The coastal zone of Brazil: an overview. J Coastal Res (Special Issue) 39:16–20 Muehe D (2003) Beach morphodynamic research in Brazil: evolution and applicability. J Coast Res (Special Issue) 35:32–42 Muehe D (2006) Erosion in the Brazilian coastal zone: an overview. J Coastal Res (Special Issue) 39:43–48 Muehe D, Corrêa CHT (1989) Beach morphodynamic and sediment transport of Massambaba beach barrier (in Portuguese). Revista Brasileira de Geociências 19(3):387–392 Muehe D, Neves C (1995) The implication of sea level rise on the Brazilian coast: a preliminary assessment. J Coast Res (Special Issue) 14:54–78 Silveira JD da (1964) Morfologia do litoral. In: AZEVEDO A de (ed) Brasil a Terra e o Homem. Companhia Editora Nacional, São Paulo, Brasil, pp 253–305 Suguio K, Martin L (1978) Quaternary marine formations of the State of São Paulo and southern Rio de Janeiro. International Symposium on Coastal Evolution in the Quaternary, I.G.C.P. Project 61, Special Publication 1, Sao Paulo Tricart J (1960) Geomorphological problems of the east coast of Brazil (in Portuguese). Bol Baiano Geogr 1:1–39

4.7 French Guiana

J. Turenne

1.  Introduction The coastline of French Guiana extends along the Atlantic Ocean between Brazil and Surinam for 370 km. It is a landscape of mangroves and salt marshes, broken here and there by outcrops of Pre-Cambrian crystalline basement (>Fig. 4.7.1). Outcrops of the Guianese Shield stand out from the coastline, and form islands in the open sea. They include the granitic islands Le Grand and Le Petit Connetable (the High and the Low Constable, 04° 50' N; 51°56' W) near the coast; Le Père (the Father, 100 m in altitude); La Mère (the Mother, rising to 111 m); L’Enfant Perdu (The Lost Child); doleritic Le Malingre (the Sickly One, rising to 55 m) and Les Iles du Salut (the Salvation Islands, 66 m) consisting of gabbros and peridotites. They also form reefs buffeted by the waves (Malmanoury Reefs). Other rocky outcrops rise from the coastal plain, which consists of marine sediment of Holocene (mainly marine clay) or Pleistocene age (fine, well-sorted sands derived from old submerged barrier sands). These include Ouanary Mountain, Montagne d’Argent (Silver Mountain, 100 m), the Cayenne Mountains (234 m, consisting of migmatites and diorites), and the granite-gneiss hills of Kourou and Organabo. The coastal region of French Guiana is a part of the large sedimentary area found in all three Guianas, and attains its maximum width in the Berbice Region (Brink­ man and Pons 1968). Clay deposits extend from the mouth of the Amazon to the mouth of the Orinoco, except for the sector between Cayenne and Organabo, where the PreCambrian basement protrudes into the sea. This tectonic partitioning, along with a slight uplift of the Pre-Cambrian basement, has influenced the morphology of the coast, the deposition of clay and the formation of sand ridges near the crystalline outcrops. There is an outer (younger) coastal plain less than 4 m above sea level, backed by an inner (older) coastal plain standing about 15 m above sea level. Sediment of the outer coastal plain are 60% clay (mainly kaolinite and some montmorillonite) and 20% fine silt, and are of Holocene age, having formed since about 8,000 years bp. They are of marine ­origin, and

include ­sediment dominated by kaolinite, which is delivered to the coast by the Amazon River and drifted alongshore to the Guianas (Dolique and Anthony 2005). The flat clay plain is varied by cheniers, ridges of coarse sand, generally oriented SE–NW, often emerging from the middle of the marine clay plain. The inner coastal plain appears to be a Late Pleistocene terrace, and consists of ridges of fine well-sorted sand, covered with savanna and forest vegetation. The Brazilian border follows the estuary of the Oyapock River, which opens to a broad gulf (Baie d’Oyapock) besides Cabo Orange, and the Approuague River opens to a similar estuary about 25 km to the west, beside Pointer Béhague. Both estuaries swing NNW, in res­­ponse to westward longshore drifting. To the west, the coastal plain is crossed by several rivers, including in westward succession the Oyac, Cayenne, Kourou, Sinnamary, Counamama, Iracoubo and Mana, all of which have small estuaries deflected to the west. The Surinam border follows the larger Maroni estuary. The coastline has prograded during the Holocene, and the river channels, banks and remnants of former river channels diversify the relief of the coastal plain. Anastomosed channels are typical of the estuaries, particularly in the Cayenne Island region, which is surrounded by such channels, as well between the Mana and Maroni rivers. The channels are vestiges of a Late Pleistocene river system, which was progressively reduced by Holocene marine sedimentation. Strong tidal currents keep the channels navigable. The remnants of a Pleistocene coastline can be observed in this region (Boyé 1963). The average range of spring tides is 3 m; that of neap tides 2.5 m. The tide wave rises through the estuaries to the waterfalls or rapids, which mark the first outcrops of the Pre-Cambrian basement. The ocean current off of the Guianas, also called the Great Southern-Equatorial Current, moves SE–NW, at between 2 and 3 knots, with a branch that runs along the coast, where sea water mingles with fresh water from the large rivers. The continental shelf, carpeted with mud and fine sand, has only a gentle slope, 0.05–1.0%; it is 92 km wide off Cabo Orange in the east, increasing to a maximum of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 4.7.1 The coast of French Guiana. A, salt marshes; B, mangroves; C, swamps.

147 km off Pointe Isère on the Surinam border, and the continental slope then drops rapidly from over 1000 m. The coastline is subject to alternations of erosion and deposition, associated with the northwestward migration of nearshore mudbanks. In a study of the port of Demerara, the Delft Hydraulics Laboratory (1962) showed that the mudbanks can form at any point on the Guianese coast as a result of refraction of incident waves, with convergence of wave energy producing zones of maximum disturbance and erosion, while divergence generates calm zones of minimum disturbance and mud deposition. The predominant

westward current along the coast drives these mudbanks from east to west, so that there is erosion and deposition as they pass along the coastline. They can be seen on satellite photographs (>Fig. 4.7.2). There is a theoretical mudbank length of 40 km, moving at 1.3 km/year, so that a mudbank arrives at 30 year intervals on any part on the coastline. Irregularities are caused by rocky outcrops on or near the coast (Turenne 1978). As the mudbanks move on, increasing wave energy brings erosion, as on the Cayenne Peninsula, a phase of erosion starting in 1963 has removed the pre-existing mangrove fringe. Landsat photographs

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⊡⊡ Fig. 4.7.2 Occurrence of mudbanks and turbid waters between Cayenne and Organabo, French Guiana: 1, erosion; 2, occurrence of turbid waters and mudbanks. Based on a comparison of 1973 and 1976 Landsat satellite photographs. N

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from 1976 show that sea waters are very muddy in the Cayenne region. The alternations of erosion and deposition is accompanied by gradual progradation of the coastline. As it advances mangroves spread seaward: a pioneer vegeta­­ tion  of Laguncularia, followed by Avicennia nitida, then Rhizophora to landward and on the banks of the estuaries. Villages such as Iracoubo, Tonate and Trou Poisson are now found inside a relatively wide fringe of mangroves. Since 1956, progradation has combined the outlet from the Iracoubo and Counamama Rivers into a single channel, and muddy deposition has greatly reduced the depth of access channels to the main ports. Cheniers are extensive on the coastal plain. They are sub-parallel ridges of coarse sand 70–200 m wide, deposited

successively by wave action on a plain of marine clay, and separated by marshy depressions, particularly on the western side of estuaries.

2.  The Coastline of French Guiana From Cap Orange (04° 24' N; 51° 34' W) to the Cayenne Peninsula (4° 50' N; 52° 22' W) the low-lying coast is bordered by mangroves. West of the Oyapock estuary the coast curves out to the Montagne d’Argent Peninsula, then mangroves are extensive to Pointe Béhague. Beyond the Approague estuary, the coast is again swampy as far as Mahury Cayenne. Island is an upland rising to more than 230 m, its north and northeast coasts cut by bays between headlands

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prolonged by rocks and islets. From Cayenne to Organabo, the coast is again low and fringed with mangroves. Low waves break over the migrating mudbanks. Locally there are sand ridges on the seaward edge of mangroves or salt marshes. At Kourou, the low coast is broken by Pointe des Roches (05° 09' N; 52°38' W) and offshore are Les Iles du Salut (05° 17' N; 52° 35' W). Further west, the coast between Organabo and Pointe Isère is fringed by salt marshes, threaded by cheniers and interrupted by tidal creeks. There are intermittent and variable sandy beaches along the shore, which provide the largest egg-laying area in the world for the luth turtle.

References Boyé M (1963) La Géologie des plaines basses, entre Organabo et le Maroni (Guyane Française). Imprimerie Nationale, Paris Brinkman R, Pons CJ (1968) A pedogeomorphological classification and Map of the Holocene sediments in the coastal plain of the three Guianas. Soil Survey Paper No. 4. Soil Survey Institute, Wageningen, the Netherlands Delft Hydraulics Laboratory (1962) Demerara coastal investigation. Delft, the Netherlands Dolique F, Anthony EJ (2005) Short-term profile changes of sandy pocket beaches affected by Amazon-derived mud, Cayenne, French Guiana. J Coastal Res 21:1195–1202 Turenne JF (1978) Sédimentologie de la Plaine Côtière, Atlas des DOM, La Guyane. ORSTOM-CNRS, Paris

4.8  Surinam     Norbert Psuty     

1.  Introduction Coastal Surinam extends for 386 km from the Marowijne River in the east to the Corantijn (Corentyne) River on the western border. The low-lying coast has been formed by waves and currents acting on the large quantities of finegrained sediment discharged by the Amazon River and the sand, silt, and clay discharged by the principal rivers (Marowijne, Suriname-Saramacca, Coppename) reaching the Surinam coast (Vann 1959). Coastal processes include waves that usually arrive from the northeast or east, producing a westward longshore drift. The tides are semidiurnal with a mean spring tide range of 2.8–3.2 m. (2.1 m at Paramaribo). The westward longshore drift is important in transporting silt, clay, and sand. Vast quantities of clay are distributed NW by the general oceanic circulation and by waves in the nearshore zone. There is some sand transport along the beach where the swash eventually strikes the coastline. In areas where sand and shell particles accumulate, coarser sediment is transported primarily by beach drifting. The Holocene coastal plain has been built up as a wedge thickening seaward during the Late Quaternary marine transgression. This recent accumulation has a subaerial extent of about 20 km at the eastern margin and gradually widens to a maximum of 140 km at the western boundary. The coastal deposits consist largely of clay with some lenses of sand and shells. On the surface, the general topography is one of low relief with several series of ­chenier ridges trending east to west and separated by wide swampy swales in which peat and mud have accumulated. The cheniers are found in groupings that extend fanshaped west from the major river mouths. Each of the rivers discharges into a narrow estuary, and there is evidence that the cheniers are composed of sand and shell derived locally from streams and the reworking of nearby ­mudbanks (Geijskes 1952). A single discontinuous chenier ridge extends along much of the present coastline, related to active beachforming processes, the availability of sandy sediment, and

the presence of wide mudbanks offshore. The cheniers are low (1.5–2.0 m), with an occasional small dune cap or hummock. In general there is a westward diminution in the volume of the sandy deposits and the number, width and height of the cheniers. Sandy beaches are present primarily in the eastern sector, where accumulations of coarser sediment are locally available, probably from the erosion of older cheniers.

2.  The Surinam Coastline West from the mouth of the Marowijne River the low swampy coast is fringed by a beach about 10 km long and 75 m wide. It has low dune hummocks approaching 0.5 m in elevation and a wide sloping sandy washover extending into mangroves on its landward margin. The Cottica River flows westward behind the swamps to join the Suriname estuary at Nieuw Amsterdam. Extensive areas have been embanked by Dutch engineers and reclaimed as agricultural polders. The coast is low and swampy west to Braams Punt, a small foreland bordering the Suriname estuary. The Suriname is joined by the Cottica and the Saramacca near the port city of Paramaibo and to the west the coastline is set back in a broad swampy lobe to the estuary of the Coppename River. To the west the coast is again set back in the province of Coronie, and has a swampy fringe, extending past the small townships of Totness and Friendship. The Nickerie River turns westward to join the Corantijn River near the port of Nieuw Nickerie.

References Geijskes D (1952) On the structure and origin of the sand ridges in the coastal zone of Surinam. Tijdschrift Koninklijk Nederlands Aardrijkskundig Genootschap 69:225–238 Vann J (1959) The geomorphology of the Guiana coast. In: Proceedings of the second coastal geography conference, Baton Rouge, Office of Naval Research Washington, DC, pp 153–187

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4.9 Guyana

Eric Bird

1.  Introduction The coast of Guyana (>Fig. 4.9.1) is about 434 km long and consists of mangroves, mudflats, cheniers, and sandy beaches fronting a swampy coastal plain. The coastal plain is dominated by the Demerara Formation (Holocene), which borders the entire Atlantic Ocean shore along this portion of the Guyana Shield. It varies from 25 to 35 km in  width and, where cleared of mangrove and drained as ­polders, is the site of extensive rice and sugar cultivation. Rivers and streams are tidal across the coastal plain. Farther inland tropical rain forest is prevalent. There are two highly variable rainy seasons, April– August and November–January, with an annual average rainfall at the coast of 2,250 mm (Brooks 1982). Humidity is high, typically 80–90% in the morning and 70–80% in the afternoon, and the average temperature 26.9°C. The continental shelf is wide, and nearshore waters shallow. Tides are semidiurnal, with a spring tide range of 2.6 m at Georgetown but less than 2 m along the open coast (1.2 m at Waini Point in the northwest). The region is in an east coast swell environment with trade monsoon influences but located south of the hurricane belt (Smith 1962). For the greater part of its length the coastline consists of mudflats, with abundant mangroves, cheniers, and occa­sional sandy beaches (>Fig.  4.9.2). Sediment from the Amazon River forms migrating mudflats and associated subtidal mudbanks (Wells and Coleman 1981). The several rivers bring some quartzose sand and large quantities of silt and clay to the coast. Waves reworking the sediment concentrate the sand fraction near high tide level, and as net sediment transport is to the west there is deposition of sandy material in that direction. Sandy beaches drift along the high tide shoreline, the quartzose fluvial sand being accompanied by shelly gravel and carbonate sand derived from shelly organisms that inhabit the intertidal and nearshore mud. Waves deposit sandy material as cheniers overlying a  muddy backshore substrate. Under the westward drift regime, the cheniers tend to form on the western side of  river mouths and trail off in a westerly direction

(Psuty  and Mizobe 2005). Their seaward margins form beaches of shells and sand, fronted by tidal mudflats, but mangroves may colonise these and form strips of swamp between successively deposited cheniers. The mangroves are of the Occidental Province, and thrive on tidal mudflats bordering estuaries and along the  coast. Avicennia is often found growing seaward of Rhizophora, an unusual pattern in the Caribbean region, probably resulting from the abundance of sand along the outer shore.

2.  The Guyana Coastline The precise location of the Surinam border along the Corentyne River estuary has still to be agreed. If the high tide line is taken as the border there is the problem of whether the mangroves and intertidal zone bordering the estuary belong to Guyana or Surinam, and where the boundary runs north across the intertidal zone. There is  also the question of the ownership of shoal areas that are partly emerged at low tide, and also migrating westward. From the Corentyne estuary the low-lying coast runs northwest to New Amsterdam. The sandy coastal fringe has a string of villages, backed by drained farmland, the  Black Bush Polder. New Amsterdam stands beside the  Berbice River estuary. The coast continues past Fort Wellington, the sandy fringe interrupted by inlets at the mouths of the Abaret, Mahaicony and Mahaica Rivers, and the polders extend to Georgetown, which stands below sea level, protected by embankments. A sea wall and promenade line the coast at Georgetown on the eastern side of the Demerara River estuary. The low-lying coast runs on to the Essequibo River estuary, the funnel-shaped mouth of which contains several alluvial islands. The river diverges on either side of Hog Island, and is then split by Wakenaam and Leguan Islands. To the north Tiger Island has the form of a barrier island running parallel to the coastline and backed by a narrow strait. The mouth of Pomeroon River has been deflected northwestward behind a sandy spit through which the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

4.9    

Guyana

⊡⊡ Fig. 4.9.1 The Guyana coast: Location map. (Courtesy Geostudies.)

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Cozier Canal has been cut. Beyond Dartmouth the coast is swampy, with extensive peat deposits, but there are sectors of sandy beach at Moruka Beach, Waini Beach, Point Kokali and Shell Beach.

Waini River also turns northwest as it approaches the coast, and flows through swamps behind the barrier spit that extends to Waini Point. Here the boundary with Venezuela crosses the coastline.

Guyana  

   4.9

⊡⊡ Fig. 4.9.2 Geomorphology of the Guyana coast: Location map. (Courtesy Geostudies.)

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References Brooks J (ed) (1982) South American handbook. Trade and Travel Publications, Bath, England Psuty NP, Mizobe C (2005) Guyana. In: Schwartz ML (ed) The Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, p 907

Smith RT (1962) British Guiana. Oxford University Press, London Wells JT, Coleman JM (1981) Periodic mudflat progradation, north­ eastern  coast of South America: a hypothesis. J Sediment Petrol 51:1069–1075

247

4.10 Venezuela

L. Ellenberg

1.  Introduction The Republic of Venezuela has a coastline about 4,993 km long (>Fig. 4.10.1). The coastal climate is subtropical, with a dry season (verano: summer) between December and April and a wet season (invierno: winter) in the rest of the  year. Maracaibo has mean monthly temperatures of 27.2°C  in January and 29.4°C in July, with an average annual rainfall of 577 mm. Tide ranges are small, with mean spring tides only locally exceeding 0.5 m. Wave action produced by the steady trade winds is from the east and northeast, with moderate energy throughout the year, and there is a strong westward longshore current. Hurricanes are very rare. Coasts exposed to strong wave action are generally cliffed, while protected coasts are often mangrove-fringed. Fringing coral reefs are found at many places, the largest

being on the eastern side of the Paraguana Peninsula. Bioerosion features are common, especially on sheltered rocky coasts. Man’s direct influence on the coast of Venezuela is still very small, less than 2% of the coastline having been altered, but indirect influences are great, notably the des­ truction of vegetation and soil cover by farming and ­grazing in recent decades, which has increased fluvial ­sediment yields, especially from the Orinoco, Unare and Tucuyo Rivers, and contributed to accelerated delta growth. The coast is unstable because of tectonic uplift of about 0.5 mm/year in the Cordillera de la Costa (Schubert et al. 1977) and subsidence in the east (Fiedler 1970). Coastal outlines are influenced by the topography of the Cordillera de la Costa, the hills of Falcon and Sucre, and the mesetas of Paraguana and Guahira.

⊡⊡ Fig. 4.10.1 The coastline of Venezuela.

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.10, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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While the variety of coastal features in Venezuela is explicable mainly in terms of geology and marine processes, climatic factors are also important (Ellenberg 1978). Those related to tropical conditions include  the  exten­­ sive coral reefs, luxuriant mangrove swamps, widespread beach rock and such features as intensive supralittoral solution on rocky shores, rapid water-layer weathering and bioerosion, and a prevalence of fine sandy beaches and alluvial shore deposition, wave quarrying being relatively unimportant (Ellenberg 1980). There is a rapid transition from humid tropical conditions in the east to warm arid climate in the west, and this is reflected in coastal geomorphology. Cliffs in soft sedimentary formations are gentle and subject to landsliding after heavy rain (>Fig. 4.10.2) in the wetter eastern sectors, and steeper and dissected only by gully erosion in the drier western sectors. Sebkhas, which are swamps in rainy seasons and saline flats in dry seasons, are found on lowlying coasts throughout Venezuela, but they are larger and more frequent in the drier west. Dunes are found only in sectors where more than 4 months of the year are dry, and where precipitation is less than 900 mm/year. However, beach rock (Reuber and Ellenberg 1979), coral reefs, and mangrove swamps show no correlation with climatic variations along the Venezuelan coast.

2.  The Coastline of Venezuela From the Guyana border the coast is low-lying and swampy, with extensive mangroves (>Fig. 4.10.1, A). Isla Cor­ocoro has sandy outer shores, and there are cheniers on the coastal plain, similar to those of Guyana. The mangroves extend around the large Orinoco delta (25,000 sq. km), where about 40 distributaries open through estuaries to a 360 km deltaic coast (Warne et al. 2002). Progradation continues, about 900 sq. km of land having been added in the twentieth century. The climate is hot and humid, and river flooding is extensive in midwinter (August–September). The mangroves are backed by swamp forests with palms, and there are some sandy areas near their mouths, as at Isla Mariusa. Longshore drifting is westward, particularly on the shores south of Trinidad, where the Caño Macareo distributary is deflected westward. Along the northwest margin of the Orinoco delta are mud volcanoes, related to rapid sedimentation accompanied by ­tectonic compression along the boundary between the South American and Caribbean plates: some are active and sparsely vegetated, others inactive and densely vegetated (Aslan et al. 2001, 2003). The swampy coast comes to an end as the rugged forested mountain Peninsula de Paria (>Fig. 4.10.1, B) runs out east towards Trinidad. This consists of strongly folded

⊡⊡ Fig. 4.10.2 Slumping cliffs at Aguide, east of Puerto Cumaredo.

Venezuela

and faulted Mesozoic limestones, dissected by deep river valleys. It is bordered by steep coasts that descend to cliffs and a beach fringe, with mangroves sparse in the bays and abundant corals on the shore, although reefs are not important. On its northern coast are many bays with sandy beaches, as at Playa Pui Puy and Playa Medina, and near the eastern end is the village of Macuro, where Columbus is thought to have landed in 1498. The coastal ridge runs westward as the more arid Peninsula de Araya on the north side of the Golfo de Cariaco, and at its western end the town of Araya is notable for its extensive salinas (salt works) in artificial lagoons which become coloured pink with Artemia, a microscopic shrimp. In the Mochima National Park steep coasts extend round the Bahia de Mochima, the Manare peninsula and the Islas Caracas, and the narrow Cerro Aceite de Castilla peninsula runs out westward beside the Golfo de Santa Fe. The Islas Caracas are barren and rocky, but have beaches and fringing coral reefs, and similar high islands continue offshore: Isla Chimana, Isla Borrachas and Isla Monos. Playa Colorado has orange sand. Puerto La Cruz is a major port and resort town with a busy seafront. West of the Golfo de Cariaco the coast becomes lowlying, with active sedimentation. A sandy coastal barrier ⊡⊡ Fig. 4.10.3 Macuto, on the steep coast of Cordillera de la Costa.

4.10

encloses lagoons on either side of the Unare delta and again at Lago de Tacarigua to the west (>Fig. 4.10.1, C). Towards Cabo Codera there are mangrove swamps in a bay sheltered from strong wave action, where sediment are trapped by the vegetation (>Fig. 4.10.1, D). Cabo Codera marks the beginning of a steep coast  (>Fig. 4.10.1, E) along the Cordillera de la Costa (>Fig. 4.10.3), which consist of strongly folded and faulted Mesozoic schists, metamorphic and intrusive rocks and Tertiary and Quaternary conglomerates. The mountains rise to the peaks of Naiguata (2,765 m) and Turimiquire (2,470 m). The steep coast descends to a narrow plain with the resort towns of Naiguata, Caraballeda and Macuto, which were severely damaged by mudslides during heavy rains in December 1999 (Perez 2001). To the west is the Parque Nacional Henri Pittier, also a steep coast with cliffs on rocky headlands, slumping slopes, beaches in small bays and mangroves in sheltered inlets. Further along the coast there are rias, such as Bahia de Turiamo. Isla Larga is fringed by beaches and reefs. Porto Cabello (>Fig. 4.10.1, F ), founded in the sixteenth century, has a sector of low-lying sandy coast that has  been  much eroded in recent times. North of Tucacas

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(>Fig. 4.10.1, G) shallow coastal water reduces wave action in the Golfo Triste, and mangrove swamps are extensive in Los Manglares (>Fig. 4.10.4) and bordering the elongated Golfete de Guare, a drowned valley south of Chichiriviche (>Fig. 4.10.5). Small islands offshore are cayos (similar to the Florida keys) with fringing reefs, ridges of coralline gravel, calcareous beaches with beach rock, and mangroves. Near San Juan de los Cayos (>Fig. 4.10.1, H) a beach ridge plain is backed by sebkhas (saline lagoons), and a fringing coral reef borders the San Juan peninsula (>Fig. 4.10.5). The reef, capped by sandy cays, extends intermittently northwest and southeast of this peninsula. EL

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Towards Puerto Cumarebo (>Fig. 4.10.1, I ) the coastal region is hilly, consisting of Oligocene to Pliocene molasse sediment, and there are steep and cliffed coastal slopes and intervening bays with recent sediment. The coast curves northwards and a narrow, low isthmus, the Istmo de Médanos (>Fig. 4.10.1, J), formed by tectonic uplift about 3,000 years ago, links Paraguana with the mainland. Its eastern shore, exposed to strong wave action, has a long beach protected by a ledge of exposed beach rock 32 km in length. The dry tabular peninsula of Paraguana (>Fig. 4.10.1, K) is a limestone area with an eastern coast exposed to strong

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⊡⊡ Fig. 4.10.4 Los Manglares.

4.10

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⊡⊡ Fig. 4.10.5 Low-lying coast around San Juan de los Cayos. CAYO DEL NOROESTE 1 km CAYO SAN JUAN

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⊡⊡ Fig. 4.10.6 Peninsula de Paraguana, a part of the Venezuelan coast that has been strongly influenced by tectonic movements. After the emergence of the Istmo de Médanos longshore drifting was diverted northward along the east coast, and a beach prograded along the eastern shore of the isthmus (a). This is now protected by a ledge of beach rock, which also extends along the southeast shore of Paraguana (b). In northeastern Paraguana (c) a fringing coral reef was raised above high tide level, and on the western coast (d) sand deposition and spit formation have ensued. West of the isthmus, in the now sheltered waters of the Golfete de Coro, the Rio Mitare (e) has built a large delta (450 sq. km). Sedimentation in and around this bay (f) has led to mangrove encroachment and the formation of saline clay plains, while the waters of the bay have become increasingly hypersaline. Dunes have moved across the southern part of the isthmus (g) and in the north of Paraguana (h) coastline progradation has occurred as the result of dunes moving from the land to the sea.

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wave action, where low cliffs alternate with beaches. A central hill, Cerro Santa Ana, rises abruptly to 830 m through well-defined vegetation zones. A fringing coral reef extends along the north-eastern coast. The western coast is more sheltered and low-lying, with high faultinduced cliffs only around Punto Fijo, which has an oil terminal and salt works. To the west the low-lying coast of eastern Zulia and western Falcón (>Fig. 4.10.1, L) is backed by alluvial deposits, dunes and Pleistocene gravel terraces. Coastal waters are shallow and wave action meagre. The longshore current is weak, especially in the east, where the bird-foot delta of the Rio Mitare has grown rapidly into shallow water ­during the past 3,000 years (>Fig.  4.10.6). Dunes, including some barchans, are migrating across the southern part of the isthmus towards the Golfete de Coro (>Fig. 4.10.7).

The narrow entrance to Lake Maracaibo (>Fig. 4.10.1, M) is fault-bounded, and consists of steep coastal slopes with almost no wave action, and mangrove fringes. Mean spring tide range increases from 0.4 m at Amuay on the west coast of the Paraguana peninsula to 1.1 m at Malecon in the approaches to Lake Maracaibo. Man­grove swamps on sheltered sectors of a shallow bay were the site of lake dwellings on piles, named Small Venice (i.e. Venezuela) when explorers arrived in 1499 (>Fig. 4.10.1, N). Lake Maracaibo (13,300 sq. km, 210 km N-S and 75 km E-W) is shallow with swampy shores, and is fed by many rivers, the largest of which is the Rio Catatumbo, which has built a swampy delta on the southwest coast. This is a region with an unusually high frequency of lightning, possibly due  to turbulence resulting from the interaction of

⊡⊡ Fig. 4.10.7 Dunes moving towards the Golfete de Coro. 69°45'W

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⊡⊡ Fig. 4.10.8 Beach ridges, mangrove swamps and sebkhas on the coastal plain of Sinamaica, Gulf of Venezuela.

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Venezuela

cold descending air from the high mountains with hot ascending air from the steamy lake. The southern part of Lake Maracaibo is fresh water, grading to brackish north towards the marine entrance, which is constricted by a long sand bar. A navigable channel 10.5 m deep was dredged in 1957, bordered by a stone breakwater. The lake occupies a Tertiary sedimentary basin which in 1917 was found to contain oil. It soon became a major oilfield, with numerous derricks standing in the water and along the shore, especially in the northeast. On the western shore of the narrow entrance strait is the port of Maracaibo, with oil refineries. West of the entrance to Lake Maracaibo (>Fig. 4.10.1, O) a low-lying coast with sandy beach ridges near Sinamaica (Tanner 1971) is backed by swamps which fringe the north and east shores of the Bahia de Uraba (>Fig. 4.10.8). The longshore current is strong on the open coast. There has been strong uplift on the Venezuelan fringe of the Peninsula de Guajira (>Fig. 4.10.1, P). Cliffs of cemented gravel up to 6 m high are being cut back by strong wave action, and there are intervening shallow bays with gravel beach ridges and dunes. Outlying Venzuelan islands include the Nueva Esparta group, of which the largest is Isla Margarita. This consists of two uplands linked by a sandy isthmus, the northern coast of which has a mangrove-fringed lagoon, La Restinga, enclosed by a sandy barrier. There are numerous beaches of white sand (Alexander and Bertness 1982). To the west is Isla La Tortuga, and farther seaward is a chain of Venezuelan islands from Islas Los Hermanos and Isla Blanquilla to Isla

4.10

Orchila, Los Roques and Isla Las Aves (Schubert and Laredo 1984). Los Roques is a small archipelago which includes a high island, Gran Roque, and a number of white sandy cays on coral reefs.

References Alexander CS, Bertness J (1982) A comparative study of modern and ancient beach morphologies: insights to the paleoclimate of Marga­ rita Island, Venezuela. J Geol 90:663–678 Aslan A, Warne AG, White WA, Guevara EH, Smyth RC, Raney JA, Gibeaut JC (2001) Mud volcanoes of the Orinoco delta, eastern Venezuela. Geomorphology 41:323–336 Aslan A, White WA, Warne AG (2003) Holocene evolution of the western Orinococ delta. Bull Geol Soc Am 115:479–498 Ellenberg L (1978) Coastal types of Venezuela – an application of coastal classifications. Z Geomorphol 22:439–456 Ellenberg L (1980) On the climatic morphology of tropical coasts (in German). Berl Geogr Stud 7:177–191 Fiedler G (1970) Seismic activity in Venezuela in association with important tectonic fault zones (in German). Geol Rundsch 59:1203–1215 Perez FL (2001) Matrix granulometry of catastrophic debris flows (December 1999) in central coastal Venezuela. Catena 45:163–183 Reuber I, Ellenberg L (1979) Beach rock in Venezuela. Acta Cient Venez 30:462–447 Schubert C, Laredo M (1984) Geology of Aves Island, Venezuela, and subsidence of Aves ridge, Caribbean sea. Mar Geol 59:305–318 Schubert C, Valastro S, Cowart JB (1977) Evidence of recent uplift on the north-central (Cordillera de la Costa) coast of Venezuela (in Spanish). Acta Cient Venez 28:363–372 Tanner WF (1971) Growth rates of Venezuelan beach ridges, Sediment Geol 6:215–220 Warne AG, Guevara EH, Asian A (2002) Late Quaternary evolution of the Orinoco delta, Venezuela. J Coast Res 18:225–253

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4.11 Caribbean Coast of Colombia

Iván Correa . Robert Morton

1.  Introduction The Caribbean coast of Colombia is about 1,600 km long, from Castilletes on the western border of Venezu­­ ela  to Cabo Tiburón at the eastern border of Panamá (> Fig. 4.11.1). It is a relatively developed area with numerous small cities and five large commerce  centres or ports (Riohacha, Santa Marta, Barranquilla, Cartagena and Turbo). Land and air access from the interior of the country is available to all medium and large cities. Colombia’s primary Caribbean island areas are the coralline archipelagos

of San Andrés, Providencia, Santa Catalina Islands and El Rosario Islands, 100 km south of Cartagena. Located at the intersection between the Nazca, Caribbean and South American plates, the Caribbean coast of Colombia is a mosaic of geologic and physiographically varied units composed of both extensive low-relief plains and medium to high relief rocky massifs. The Caribbean coastal zone is crossed by several active faults that define its main morphostructural units. The area has been classified as an intermediate seismic risk zone.

⊡⊡ Fig. 4.11.1 Caribbean coast of Colombia.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Caribbean Coast of Colombia

The climate of the Caribbean coast (IGAC 2002) depends on the annual displacement of the Intertropical Convergence Zone and on particular orographic influences of the Sierra Nevada de Santa Marta massif. There are generally two rainy periods (April–May and October– November) and two dry periods (December–April and July–September). Maximum annual precipitation for the Colombian Caribbean does not exceed 2,500 mm. Minimum values are within the desert region of the Guajira Peninsula (yearly mean about 267 mm), and maximum values are at the Sierra Nevada de Santa Marta massif (yearly mean 2,000 mm). Mean air temperatures for the Caribbean coast are less than 24°C. Tides along the Caribbean coast are of mixed semidiurnal type, with maximum amplitudes of 60 cm. Trade winds predominate, mainly from the east, north and northwest, at the Guajira Peninsula, and from the NE to NW, south of the Sierra Nevada de Santa Marta. Net longshore drift along the Caribbean coast of Colombia has a dominant southward component, minor reversals to the northeast occurring during the rainy periods when south winds become dominant in some sectors. The Caribbean coast of Colombia is out of the zone of direct influence of tropical cyclones, but is affected by perimeter influences, especially along the northern Caribbean, between La Guajira and Cartagena, where high waves cause extensive shore erosion and lowland flooding.

2.  Southern Caribbean Coast The southern Caribbean coast of Colombia extends for about 550 km from the Barú Peninsula to the Gulf of Urabá (> Fig. 4.11.1). Except for the Gulf of Morrosquillo, the region is cut into sedimentary rocks of the Sinú Belt, a stratigraphic unit composed by Oligocene to Pliocene sequences of turbidities, hemipelagic and terrigenous marine deposits, strongly affected by differential tectonic movement driven in part by mud diapirism and associated phenomena (Duque-Caro 1984; Vernette et al. 1992). Along the littoral fringe, rocks of the Sinú belt are mainly claystones, mudstones and conglomerates. Coral reef limestones outcrop on the tops of coastal hills, which are up to 150 m high. Offshore and inland mud diapirs and mud volcanoes are common on the Sinú Belt, some of them with violent historical eruptions that modified the bathymetry and surface topography of the order of metres (Vernette et al. 1992). Quaternary deposits of the Sinú belt coast are mainly sandy and muddy deltas (Sinú-Tinajones deltaic complex,

Turbo and Atrato deltas) and alluvial valley fills. Along the coast, there are also Holocene terraces up to 36 m high, best developed south of the Gulf of Morrosquillo. Holocene reefs on the shallow platform of the Sinú Belt are located on positive-relief bottom features formed by mud ­diapirism (Duque-Caro 1984; Vernette et al. 1992) and are most common between Cartagena and the Gulf of Morrosquillo (El Rosario Islands). The Gulf of Morrosquillo, 100 km south of Cartagena, is a greatly modified, low wave-energy environment. Internal lagoons and mangrove swamps are fronted by protective beaches and extensive reclamation projects. Southwest of the Gulf of Morrosquillo, the coastal plain is composed of extensive detrital deposits of the Holocene Sinú River delta. The main entrance to the Bay of Cispatá changed abruptly in 1938 as the Sinu River enlarged an irrigation channel cut at Tinajones sector (Serrano 2004; Correa et al. 2007). As a result of channel widening, the Cispatá bay area became depleted of sediment and was submerged by the sea. More than 10,000 ha of cultivated rice land was lost at its margins. The lobate, symmetrical Tinajones delta has an approximate area of 26 km2; both flanks are subjected to intense beach erosion. From Punta La Rada to the Gulf of Urabá, the southern Caribbean coast has an S 50° W trend and is dominated by cliffs cut into deformed sequences of Tertiary mudstone and claystone. The cliffs, which are up to 36 m high, form the seaward limit of late Holocene wave-cut and depositional marine terraces. The coast exhibits a serrated pattern controlled by hard rocky headlands coinciding with structural axes. Shore retreat is the dominant historical trend along most of this coastline. A recent example of land loss along the southern Caribbean coast is in the Arboletes-Punta Rey area. Here the effects of natural shore and cliff erosion are exacerbated by intensive beach-sand extraction to build up the village of Arboletes. The natural processes and human activities caused complete erosion of the 1.5 km long Punta Rey Peninsula between 1957 and 2000 (Correa et al. 2007). Elimination of the wave protection provided by the peninsula triggered extensive cliff and beach erosion to the south. Accelerated erosion also affected the Arboletes mud volcano, the main tourist attraction of the city. Currently stable shores along the southern Caribbean coast are found only at few specific, unmodified stretches that receive relatively high sand supply. The Gulf of Urabá is a relatively wave protected environment with a general N-S orientation and a maximum width of 20 km at its northern end. On its NE part the Gulf shores have narrow beaches backed by cliffs cut into ­sedimentary sequences of the Sinú Belt. The southeastern

Caribbean Coast of Colombia

coast of the Gulf is dominated by erosion, some local deposition associated with recent human activities – the Turbo River delta being the most important. The SW side of the Gulf of Urabá is defined by the prograding Atrato River delta, which has 16 active distributaries protruding into the 25 m deep Gulf. The Atrato River is 730 km long and drains an area of about 36,000 km2 of the western Colombian Cordillera and the Serranía del Baudó, on the Pacific Coast. Rainfall in its drainage basin is 8 m/yr, which supplies a mean water discharge of 2,700 m3/s and a sediment yield of about 11 × 106 ton/year (Restrepo and Kjerfve 2000). Northwest of the Gulf of Urabá and the Atrato River delta, the Caribbean coast changes to basaltic cliffs alternating with wide alluvial valleys and beach ridges. Absence of sediment from the Atrato River, permit the development of fringing coral reefs which become progressively more abundant toward Panamá.

3.  Northern Caribbean Coast The northern Caribbean coast extends for about 600 km, between Castilletes and the city of Cartagena. At its northern sector, the Guajira Peninsula is a main morphostructural element consisting of tectonically raised blocks of metamorphic, granitic and sedimentary rocks (Jurassic to Tertiary in age), adjacent to sedimentary basins and

⊡⊡ Fig. 4.11.2 Aerial view of Cabo de la Vela, showing 5–10 m high cliffs cut into terraces of serpentinitic rocks covered by small dunes and a veneer of wind-transported sand. (Courtesy Diego Zapata.)

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g­ rabens filled with Tertiary limestones, cherts and claystones. The Quaternary of the Guajira Peninsula is mainly represented by extensive colluvial and alluvial deposits, and recent sandy barriers and marine lagoons. The Guajira Peninsula coastline from Castilletes (Gulf of Maracaibo) to Dibulla is about 280 km long. Spits, bars and lagoons predominate along the internal shore of the Gulf of Maracaibo, whereas narrow beaches and cliffs are dominant along the Bahia Honda-Cabo de la Vela shore (> Fig. 4.11.2). South of Cabo de La Vela, the Guajira Peninsula coast is dominated by transgressive narrow beaches, minor ­deltaic accumulations and spit-lagoon segments near the mouths of the primary coastal rivers. These coastal-plain deposits are located seaward of extensive erosional platforms cut into Tertiary mudstones. South of the Guajira Peninsula, the Sierra Nevada de Santa Marta massif is the highest coastal mountain in the world, reaching 5,800 m at Pico Bolívar, 60 km from the coastline. It is composed of Cretaceous metamorphic rocks (predominantly schists and gneisses) and quartz-dioritic intrusive rocks of Tertiary age. The Quaternary of the Sierra Nevada de Santa Marta is mainly represented by colluvial and alluvial valley fill and by recent beach deposits. The coastlines of the northwestern part of the Sierra Nevada de Santa Marta are indented, reflecting the alternation of rocky headlands and deep, NNW trending ­tectonically controlled bays. The headlands are typically

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100–150 m high plunging cliffs cut into schists and granites, whereas the embayments front alluvial valley deposits. The bay shores are typically steep, reflective beaches composed of very coarse sand to granules eroded from adjacent cliffs and alluvial deposits. Wide pocket beaches and tombolos are common along sectors with abundant rocky erosional remnants and stacks (> Fig. 4.11.3). West of the Santa Marta massif, between Santa Marta and Barranquilla, the coastal zone includes the extensive Ciénaga de Santa Marta shallow ( Fig. 4.11.4). Most cliffs are fronted by narrow beaches supplied by erosion of Tertiary formations and sand eroded from the Magdalena River delta. Near Cartagena, the coastline is dominated by depositional features fronting inactive cliffs, mainly sandy spits and bars that modulate the coastal indentations. Historical information on this area (Vernette et al. 1992) shows large shoreline changes associated with mud diapirism along this segment of coast, including erosion of a 12 km long, E-W oriented spit reported in 1794 and the formation of up to 7 km2 tombolos. Extensive land reclamation has been carried out to enlarge the city of Cartagena. The Tierrabomba Island and the Barú Península are composed of Plio-Pleistocene mudstones and limestones that show evidence of recent relative sea level changes, including wave-cut platforms, elevated caves and stacks.

⊡⊡ Fig. 4.11.3 View southward over Cañaverales Bay at the Santa Marta massif, showing highly reflective coarse sand beaches, tombolos and a rocky NW coastline.

Caribbean Coast of Colombia

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⊡⊡ Fig. 4.11.4 Steep cliffs cut into Tertiary sandstones and mudstones and fronted by narrow beaches at El Morro Point, south of Barranquilla. (Courtesy Jesus A. Pérez.)

The features are as much as 5 m above present sea level and have been dated as late Holocene. Steep erosional cliffs and narrow calcareous pocket beaches dominate the outer coastlines of Tierrabomba Island and the Barú Peninsula. Extensive mangrove swamps are found along the margins of the Cartagena and Barbacoas Bays.

References Correa ID, Acosta S, Bedoya G (2007) Análisis de las causas y monitoreo de la erosión litoral en el departamento de Córdoba. Fondo Editorial Universidad EAFIT, Medellin, pp 1–127

Duque-Caro H (1984) Estilo estructural, diapirismo y episodios de acrecimiento del terreno Sinú – San Jacinto en el Noroccidente de Colombia. Boletín Geológico Ingeominas 27:1–29 IGAC (2002) Atlas de Colombia 5th edn. Bogotá, p 342 Restrepo JD, Kjerfve B (2000) Magdalena River: interannual variability (1975–1995) and revised water discharge and sediment load estimates. J Hydrol 235:137–149 Serrano BE (2004) The Sinú River delta on the northwestern Caribbean coast of Colombia: bay infilling associated with delta development. J South Am Earth Sci 16:623–631 Vernette G, Maufret A, Briceño L, Gayet J (1992) Mud diapirism, fan sedimentation and strike-slip faulting, Caribbean Colombian margin. Tectonophysics 202:335–349

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4.12 Caribbean Coasts, Panama to Belize

Anja Scheffers · Tony Browne

1.  Introduction The Caribbean coasts of Central America extend for about 2,200 km from the Atrato lowlands in northwestern Colombia to Tehuantepec in Mexico. The geology of Central America is complex. Two large areas of different tectonic history and geological structure have been recognized: namely the Isthmian Link from NW Colombia to southern Nicaragua and Nuclear Central America, from southern Nicaragua to SE Mexico. The Isthmian Link (Weyl 1980; Dengo 1985) has a basement of Late Jurassic to Late Cretaceous age, which consists mainly of basalts with associated radiolarites and clastic sedimentary rocks. The basement is overlain by thick volcano-clastic rocks of Late Cretaceous to Pliocene age, with some interbedded limestones. The prevailing structures are parallel to the geographical shape of the isthmus. South of the mountains of Guatemala the metamorphic basement is covered by Mesozoic sedimentary rocks: Upper Triassic (in Honduras), Jurassic sediment and Cretaceous carbonates and red beds, intruded by granitic bodies. These, in turn, are overlain in extensive areas by Tertiary red beds and ignimbrite plateaus. A Lower Palaeozoic metamorphic and igneous basement (probably including some Pre-Cambrian rocks) forms some of the high mountains in northern Nicaragua, Honduras, Central Guatemala and SE Mexico. North of the mountains of Guatemala, extending westward into Mexico, the basement is overlain by a thick sequence of upper Palaeozoic sedimentary rocks, which in turn are covered by Upper Jurassic and thick Cretaceous to Eocene strata. Upper Tertiary marine sediment are restricted to small areas. The geological structure north of the Central Cordillera is dominated by a Laramidian fold belt that extends from SE Mexico eastward across Guatemala and southern Belize, forming an open arc concave to the north and with an ENE trend as it reaches the Caribbean coast. North of the fold belt the structures vary from gentle to nearly horizontal in the large area of northern Guatemala and Belize and the Yucatan Peninsula in Mexico.

A common feature of Central America is a Quaternary volcanic mountain chain that runs NW–SE, and is superposed on the older structures. The continental divide in Central America is generally closer to the Pacific, so that the longer streams drain to the Caribbean Sea. A high rainfall (5–6 m/year), notably during hurricanes, yields a large river discharge and vast amounts of fine grained sediment are transported to the Caribbean coastline, where they are distributed alongshore and mostly form coastal lowlands with beaches and beach ridges. In Panama and Costa Rica the continental divide is more central, except in the area east of the Panama Canal, where it is very close the northern coastline and results in short steep streams flowing down to the Caribbean. The coast is microtidal (around 0.5 m spring tide range). The Caribbean coast is divided into four sectors according to their similarities in morphology and origin: Cape Tiburón on the Colombian border to Port Limon in Costa Rica, Port Limon to Cape Gracias a Dios (NicaraguaHonduras), Cape Gracias a Dios to the Sarstun River mouth (Guatemala) and from there to Chetumal Bay on the Mexican border. The geomorphology and ecology of this coast has been described by Yanez-Arancibia (2005).

2.  Panama The state of Panama has the longest Caribbean coastline of Central America, more than 600 km. West from Cape Tiburón on the Colombia border the mountains are close to the coast and the beaches are interrupted by frequent long cliffs, with some fringing coral reefs. The coast is bordered by numerous small islands (the Archipelago de San Blas), made up by a drowned part of the basic igneous rock basement of southern Central America. Towards the Canal entrance there are recently uplifted shore platforms. In many places the tropical forest almost reaches the beach. From Manzanillo Point to Chiriqui Point the north-facing coast is formed by a succession of narrow beaches separated by low cliffs where the mountains reach the sea. The coastline becomes ­intricate behind the Peninsula

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.12, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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de Valiente, where the Lagoa de Chirigui and Almirante Bay are backed by Tertiary sedimentary and volcanic rocks, which also form the islands inside the bay. A part of the northern bay shows partly drowned cockpit karst.

3.  Costa Rica The coastline of the Caribbean side of Coast Rica is much shorter than that of the Pacific side, only about 200 km in length. Offshore the Miskito Keys and banks are formed by coral reefs, while the islands of Great and Little Corn consist of volcanic rocks and present small beaches separated by cliffs. At Mona Point, Cahuita Point and Port Limon there are uplifted Pliocene coral reefs and Holocene fringing reefs. North of Point Limon the coastal plain widens as the mountainous hinterland recedes. There are long gently curving beaches, some with concentrations of magnetite and ilmenite. To the NW long sandy barriers are backed by narrow coastal lagoons, notably the Tortuguero Lagoon. The beach sands are derived largely from Quaternary volcanic rocks (Nieuwenhuyse and Krooneneberg 1994). The barriers continue to the San Juan River delta.

4.  Nicaragua The Caribbean coastline of Nicaragua runs north for more than 450 km. San Juan River, which drains the Nicaragua lake basin, forms the boundary with Costa Rica. To the north is the Costa de Mosquitos between Cabo Gracias a Dios and Bluefields where mangroves, notably Rhizophora mangle, are extensive around lagoons and on swampy shores. Coastal sediment are derived mostly from low cliffs of Tertiary volcanic rock, but they include material from the metamorphic basement and Cretaceous limestones. There are several elongated lagoons bordered by sand barriers, the largest being the Perlas Lagoon. In the northern part of the coast, near Puerto Cabezas, extensive Plio-Pleistocene low terraces occur, formed largely of quartzose sand and pebbles. The coastal plain narrows towards Cape Gracias a Dios, at the mouth of the Coco River, which forms a prograding delta.

5.  Honduras The Caribbean coast of Honduras is at least 600 km long, excluding large islands (up to 50 km long) in the north.

From Cape Gracios a Dios the coast turns NW and a long sandy barrier encloses the large Caratasca Lagoon (Helbig 1959). This lagoon includes several small sandy islands, and is surrounded by a low swampy area with extensive mangroves. It was probably formed as a result of changes in the lower course of the Patuca River leaving a broad embayment across which a barrier was built by sediment delivered by the Coco River and transported alongshore to the northwest. The hinterland of the lagoon has extensive, low Plio-Pleistocene terraces. West of the Patuca delta the mountain ranges locally reach the coast. Between Cape Camaron and Port Cortes long narrow beaches extend between cliffs of Palaeozoic metamorphic rock and Tertiary ignimbrites, except near the mouths of the Aguan and Ulua Rivers, where the valleys merge into wider coastal plains, some swampy and with mangroves. The large rivers have provided a considerable amount of sediment, including sand that has been built by storm waves into beach ridges. Longshore drifting is westward, and has led to the growth of sand spits, such as the spits that border Brus Lagoon near the mouth of Patuca Ruver and Cape Camaron. Sand drifting west from the mouth of Aguan River has built a spit that partly encloses a bay at Port Trujillo. The beach sands are heterogeneous because the sediment source consists of metamorphic rocks, limestones, granitic intrusive rocks, and ignimbrites. Some of the sands have a large quartz content derived from the granitic rocks and the rhyolitic ignimbrites. In 1961 Hurricane Hattie caused extensive coastal flooding in British Honduras. damaging coral reefs and eroding and destroying sand cays (Stoddart 1962). The Bay Islands (Islas de la Bahia), north of the Honduras coast, are a continuation of the mountain system of Nuclear Central America. They are formed mostly by metamorphic rocks and one small volcanic cone (Utila). Their beaches are small and discontinuous, separated by low cliffs, and the whole island group is surrounded by a submarine coral reef.

6.  Guatemala The Sarstun River flows into Amatique Bay, which extends about 100 km along the Guatemalan coast. This bay is located on a graben, extending inland past Lake Izabal and continuing seaward as the Cayman Trough. The bay is separated from the sea by a large sand spit, Manabique Point, formed of sediment delivered by the Motagua River and transported NW by longshore drifting. The inner side of the bay is a low hilly area formed by Miocene limestones

Caribbean Coasts, Panama to Belize

and cut by the Rio Dulce gorge, the outlet from Lake Izabal. This sector has only small discontinuous beaches. The Motagua River delta is partly covered with mangroves. The beach sands include pumice carried by the Motagua River from the Quaternary volcanic deposits of the Guatemalan highlands.

7.  Belize The Belize coast has an overall length of about 300 km with some additional islands on the outer reef. The coast is dominated by carbonate deposition and reef growth (Stoddart et al. 1982; Gibson and Carter 2002). Some beaches  have been cemented by carbonate precipitation (Gischler and Lomando 1997). One of the largest barrier reefs in the world (more than 200 km long, with additional coral islands outside and atolls inside), second only to Australia’s Great Barrier Reef, forms a prominent feature in this area. It extends from the Seal Cays in southern Belize to Bahia de la Asencion in Mexico, and is interrupted by several gaps, notably a large one off the mouth of Belize River. Hurricane impacts on reefs were reported by Stoddart (1962, 1965). The reef is separated from the main coast by a shallow shelf filled with carbonate sediment (Bonis et al. 1970). The coastline and the distribution of the offshore coral reefs are structurally controlled by a series of sub-parallel normal faults trending NNE. Oil exploration drilling and marine seismic profiling have shown that the major reefs, such as Chinchorro, Turneffe, Ambergris, and Key Long, have been built since the Pliocene over basement structural highs, while the smaller linear ones are over the upthrown edges of tilted fault blocks. North from the Guatemalan border Cretaceous limestones outcrop along the coast, forming prominent cliffs separated by short, narrow beaches near Punta Gorda. North to Stann Creek Palaeozoic rocks of the Maya Moun­ tains come close to the coast, and the beach sand contains much quartz. North of Stann Creek the coastal plain widens and there are elongated lagoons behind sandy barriers.

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The low-lying area around Belize City is formed by fine calcareous sediment deposited by the Mopan River, which forms a small delta on which the city was built. The coastline of northern Belize is formed by a series of calcareous sand barriers enclosing lagoons with mangrove shores, bordered inland by Plio-Pleistocene terraces, also formed predominantly of limestone fragments derived from the older carbonate rocks (Wright et al. 1959). Luxuriant mangroves are found in the shelter of the  barrier reef on carbonate platforms elevated upon mean sea level. Many of them show distinctive species ­patterns, depending on the distance to open water and the  duration of submergence, related to the level of the substrate.

References Bonis S, Bohnenberger O, Dengo G (1970) Mapa Geologico de la Republica de Guatemala (1:500,000). Instituto Geografico Nacional, Guatemala Dengo G (1985) Caribbean Central America. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 117–124 Gibson J, Carter J (2002) The reefs of Belize. In: Cortes J (ed) Latin American Coral Reefs. Elsevier, Amsterdam Gischler E, Lomando AJ (1997) Holocene cemented beach deposits in Belize. J Sediment Petrol 110:277–297 Helbig K (1959) Die Landschaften von Nordost Honduras, Geographische Kartographische Anstalt Gotha Nieuwenhuyse A, Kroonenberg SB (1994) Volcanic origin of Holocene beach ridges along the Caribbean coast of Costa Rica. Mar Geol 1201–1202:13–26 Stoddart DR (1962) Catastrophic storm effects on the British Honduras reefs and cays. Nature 196:512–515 Stoddart DR (1965) Re-Survey of Hurricane Effects on the British Honduras Reefs and Cays. Nature 207:589–592 Stoddart DR, Fosberg FR, Spellman DL (1982) Cays of the Belize Reef and Lagoon. Atoll Res Bull 256:1–76 Weyl R (1980) Geology of Central America. Bornträger, Berlin Wright ACS, Romney DH, Arbuckle RH, Vial VE (1959) Land in British Honduras. Colonial Research Publications, London Yanez-Arancibia A (2005) Middle America, Coastal Ecology and Geo­ morphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science, Springer, pp 639–645

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4.13  Mexico, Caribbean Coast    Mario Gutiérrez-Estrada · Mario Arturo Ortiz-Perez     

1.  Introduction The Caribbean coast of Mexico is about 2,300 km long. The NE trade winds blow onshore from the Caribbean and produce waves on the Mexican coast. There are occasional hurricanes. Tide ranges are smaller than on the Pacific coast. On the Gulf Coast Campeche has a mean tide range of 0.3 m, Veracruz 0.5 m, Tuxpan 0.9 m and Tampico 0.4 m. The Caribbean coast is classified as a marginal sea (Inman and Nordstrom 1971; Carranza et al. 1975), and includes the low-lying Yucatan Peninsula. This passes westward to a coastal plain with broad Quaternary alluvial terraces, which varies in width from almost zero where the San Andres Tuxtla and Jalapa highlands reach the coast to about 150 km near the major rivers. It is backed by the Sierra Madre foothills. The continental shelf is wide off Tabasco out to the Campeche Banks, and narrows from Veracruz to Tampico. The eastern shore of the gently rolling, karstic Yucatan Peninsula is irregular and indented by bays backed by extensive swamps. The western shore is rather simple with minor irregularities and narrow marshes. There are extensive lagoons and barriers across the northern, eastern and western sides of the emerged peninsula. The coast declines to a plain towards the Laguna de Terminos. Towards the Tonala River are coastal marshes, swamps, and shallow lagoons (Murray 1961; Gutierrez-Estrada 1982). It is an active depositional coast. Inland there are deeply dissected Quaternary river and marine terraces. The Tabasco plain is a coastal lowland with swamps, marshes, and lakes along the coast westward from the Laguna de Terminos to the Tonala River (Murray 1961; Psuty 1966; Thom 1967). The largest lowland area is in eastern Tabasco, where the Grijalva and Usumacinta rivers have built alluvial plains. The coast consists of lagoons and barriers with sand hills and dunes. The Tabasco segment includes parts of Campeche, Tabasco and Veracruz. The coastal plain averages 90 km in width; it is narrowest in western Tabasco and eastern Veracruz, where there are

elevations of around 300 m. It widens into the basin of Coatzacoalcos River. The Isthmian embayment, from Yucatan to the Jalapa highlands, is divided into the Tabasco and Veracruz segments at the Sierra de Los Tuxtlas. Sand dunes extend from the Sierra de Los Tuxtlas to the Jalapa highlands, and are well developed between Alvarado and Veracruz, where they form the barrier between the Laguna de Alvarado and the Gulf. There is a coastal convex arc, apparently related to volcanism and uplift, and the coastline is fairly regular but for several cuspate bars and lagoons. The plain gradually narrows northward from the Papaloapan River basin. The northern part of the Veracruz segment is developed principally on volcanic material of late Tertiary and Quaternary ages. Maximum elevations are of the order of 200 m. Coral reefs are common in the vicinity of Veracruz city. North of Veracruz the coastline consists of shallow lagoons and marshes partially enclosed by spits and barrier islands. Laguna de Tamiahua and associated cuspate bars and reefs are the principal irregularities. North to the Panuco River Tertiary hills have elevations of the order of 100 m just west of Laguna de Tamiahua (> Fig. 4.13.1). Beyond the Panuco River the coastal plain is interrupted by the Sierras de Tamaulipas, Aldama. and Madre Oriental. The Tampico embayment is a physiographic unit of the coastal plain in east-central Mexico. The plain slopes gently eastward from the Sierra Madre Oriental to the Gulf, and has a maximum width of about 150 km. There are flat areas in its central part and along the Pnuco, Tamesi, and Tuxpan rivers, where marshes and swamps are extensive. Inland are the uplifted ranges of the Sierra Madre Oriental, from which gravelly fans and terraces extend towards the coast. North of Tampico the coastal plain narrows in front of the Sierra de Tamaulipas, then widens where the Laguna Madre is fronted by a long coastal barrier breached by a few marine inlets. Marshes and mangroves are common. The coastline curves gently NNE towards the lobate delta of the Rio Grande (Bravo) on the Texas border.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_4.13, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Mexico, Caribbean Coast

⊡⊡ Fig. 4.13.1 Predominant coastal landforms of the Caribbean coast of Mexico. (Courtesy Geostudies.) Bravo R.

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Carranza-Edwards A, Gutierrez-Estrada AM, Rodriguez-Torres R (1975) Morphotectonic units of the Mexican coast (in Spanish). An Centro Cienc del Mar y Limnologia, Univ Nal Auton Mexico 2:81–88 Gutierrez-Estrada M (1982) Geomorphology and Recent sediments of the Atasa-Pom lagoon system, Campeche, Mexico (in Spanish). An Centro Cienc del Mary Limnologia, Univ Nal Auton Mexico 9:89–100

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Inman DL, Nordstrom CE (1971) On the tectonic and morphological classification of coasts. J Geol 79:1–21 Murray GE (1961) Geology of the Atlantic and Gulf coastal province of North America. Harper, New York Psuty NP (1966) The geomorphology of beach ridges in Tabasco, Mexico. Louisiana State University Coastal Studies Institute, Technical Report 30:1–51 Thom BG (1967) Mangrove ecology and deltaic geomorphology: Tabasco, Mexico. J Ecol 55:301–343

5.0 Caribbean Islands – Editorial Introduction

The Caribbean islands (Blume 1974) lie in a broad arc, starting east of the low-lying Mexican peninsula of Yucatan with > Cuba, which has extensive lowlands. The Greater Antilles include three more large and mountainous islands: > Jamaica, > Hispaniola and > Puerto Rico extending eastward. There is then a long curving island chain, the Lesser Antilles, which continues as a series of high volcanic islands southward from Anguilla to Grenada, with an outer chain of low-lying islands that includes Barbuda, Antigua, Guadeloupe and Barbados. They then extend westward as the Dutch Lesser Antilles off the Venezuelan coast. The islands formed along an arcuate zone of crustal weakness between the North and South American crustal plates. The inner Lesser Antilles originated as a chain of submarine volcanoes, several of which grew above sea level. Volcanic eruptions have continued, notably at Souifriere on St. Vincent and Chances Peak on Monserrat, and there are also hot springs and solfatara. The outer chain of the Lesser Antilles consists of islands of limestone on old volcanic foundations. Beaches on volcanic islands are mainly of dark basaltic sand, often with gravel and boulders. Coral reefs (particularly fringing reefs) are extensive and beaches derived from these are of pale coralline sand and gravel. White ­calcareous beach sands have also been derived from shelly

debris, algae and other sea floor organisms. Beach rock, formed by cementation of sand by carbonates, has been exposed by erosion of some beaches. The outer shores of the Lesser Antilles north from Grenada and Barbados receive ocean swell from the Atlantic, which also breaks on the northern shores of Puerto Rico, Hispaniola, and to a lesser extent (because of protection by the reefs, banks and islands of the Bahamas to the north) Cuba. The swell is accompanied by waves generated by the NE trade winds (with some from the SE in summer) and these are dominant within the Caribbean Sea, shaping beaches on the southern coasts of Puerto Rico, Hispaniola and Jamaica. During summer and autumn hurricanes occur, moving in from the east over the Lesser Antilles, through the Caribbean Sea, across the Greater Antilles, and into the Gulf of Mexico. They cause erosion of beaches and cliffs, and damage to coral reefs and coastal vegetation. Mean spring tide ranges are small, generally less than 0.5 m. Mangroves occur locally in sheltered situations such as reef lagoons, inlets and estuaries and seagrass beds are extensive in shallow waters.

Reference Blume H (1974) The Caribbean islands. Longman, London

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

5.1 Cuba

Ridel Rodríguez

1.  Introduction The island of Cuba is about 1,250 km long and 30–190 km wide, and has an irregular coastline about 5.800 km long (> Fig.  5.1.1). The Cuban archipelago has four island groups: the Los Colorados archipelago, the SabanaCamagüey (Jardínes del Rey) archipelago, the Jardínes de la Reina (Gardens of the Queen) archipelago and the Los Canarreos archipelago. Cuba is geologically complex, and forms the northern part of the Antilles island arc (Khudoley 1967; Fairbridge 1975), with Jurassic and Cretaceous rocks, mainly in the mountainous areas, and Palaeogene to Quaternary in the bordering lowlands. About two-thirds of the island consists of limestone, and karstic features are well developed, particularly in the eastern section. The long and narrow island of Cuba has generally short and steep rivers. The longest is the Cauto River (370 km), followed by the Sagua la Grande River (163 km) and Zaza River (155 km). Lying immediately south of the Tropic of Cancer, Cuba has a humid to subhumid tropical climate, with two clearly

defined seasons, the dry one (winter) from November to April, and the rainy one (summer) from May to October, with continuous breezes the whole year. The north coast is exposed to the NE trade winds which bring heavy rainfall in the summer months: Havana has a mean annual rainfall of 1,226 mm, of which 868 mm fall between May and October. Mean annual temperature is 24°C (20°C in winter and 27°C in summer). In general rainfall is more abundant in the west than in the east of the country. The south coast is somewhat drier, but the whole of the island coast is subject to occasional hurricanes, especially in August, September and October. Tide ranges are small: Havana has a mean spring tide range of only 0.5 m.

2.  The Cuban Coastline The Cuban coastline is extremely irregular and intersected by numerous gulf and bays, such as the Batabanó and Guacanayabo Gulfs. There are good natural harbours in the bays. Coral reefs and keys extend along the coast. Parts

⊡⊡ Fig. 5.1.1 Cuba. Location map. (Courtesy Geostudies.)

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of the coast are steep and cliffed; others have beaches or mangroves. Although the marine provinces bordering Cuba are rich in biogenic sediment, beaches on the more exposed parts of the coast have been eroding in recent decades. 86% of Cuban beaches are affected by erosion (>Fig. 5.1.2), although on the NE coast this drops to 58.6% (Juanes 1996, Rodríguez and CÓrdova 2006). Erosion is due to a deficit in sand supply, a rising sea level, the effects of hurricanes and human impacts. Progradation has continued where there is a sand supply: some river sediment yields have been augmented by the effects of open-cast nickel mining in their catchments, and this has led to local accretion near river mouths. Havana stands beside a coastal inlet formed by Late Quaternary marine submergence of a valley mouth, but there are terraces 8 and 16 m above sea level indicating Pleistocene uplift, and others ranging laterally between 15 and 33 m, and 25–51 m, indicating transverse tilting (Ducloz 1963). West of Havana as far as Punta Gobernadora the coast is fringed by low cliffs and sandy beaches with inlet and bays such as Mariel Bay. Mangroves are extensive along the NW coast (>Fig. 5.1.3), mainly behind sandy barriers and spits or fringing coral reefs, which develop west of Punta Diamante and are close inshore off Punta Gobernadora.

In  the hinterland are the Pinar del Rio Mountains, from which short, steep rivers descend to coastal inlets. The chain of coral reefs and cays becomes the Archipielago de los Colorados, sheltering the coast round to Golfo de Guanahacabibes at the western end of the island. The shores have a mangrove fringe, particularly in the Bahia de Guadiana. The limestone Peninsula de Guanahacabibes has mangroves on its northern shores, but the southern coast has cliffs and bluffs, and is beachfringed eastward from Cabo de San Antonio past Cabo Corrientes to Cabo Francés on a coast that faces deep water and receives relatively strong wave action. East of Cabo Francés there are sectors of mangrove and some enclosed lagoons: Laguna de Cortez, Laguna Cheve. There are minor headlands between beaches (playas) edging a coastal plain, Llanura Aluvial del Sur. The continental shelf widens eastward across the ­relatively shallow Golfo de Batabanó, and Cayos de San Felipe and extends towards a large island to the south, Isla de Pinos. This has hilly ranges of metamorphic rock amid broad lowlands. Nueva Gerona stands on the north coast, and the western and eastern shores are embayed, while the curved southern coast facing deep water is partly cliffed, and passes eastward into a chain of reefs and small islands - Archipielago de los Canarreos, curving east towards Cayo Largo.

⊡⊡ Fig. 5.1.2 Erosion at Guardalavaca Beach, caused by storm waves.

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5.1

⊡⊡ Fig. 5.1.3 Mangroves near Zapata.

The northern coast of the Golfo de Batabanó has sandy beaches and mangrove-edged bays, but becomes swampy as wave energy diminishes eastward into the Ensenada de  la Broa, an embayment on the northern side of the Peninsula de Zapata. This low-lying limestone peninsula has swampy northern shores and an irregular cliffy southern coast. To the south–east a deep and narrow submarine canyon, the Golfo de Cazones, curves out southeastward across the narrowing continental shelf. A wide slot-shaped marine inlet, Bahia de Cochinos (Bay of Pigs) is set into the coast to the east, and beyond that is the Bahia de Cienfuegos, with a narrow entrance strait. To the southeast the coast becomes steep below the Sierra del Escambray, a range of mountains rising to Pico San Juan (1,140 m), but remains beach-fringed. Next comes a coastal lowland, with mangroves behind small barriers and spits, running out to the beach-fringed Peninsula de Ancon (> Fig. 5.1.4). At Trinidad a recurved spit, Punta Casilda, marks the beginning of a coast less exposed to strong wave action as the continental shelf widens again eastward, with a chain of coral reefs and cays near its outer edge forming the Archipielago de Jardines de la Reina. The mainland coast remains low-lying, with fringing mangrove swamps, notably around the shallow Golfo de Ana Maria. The River Najasa flows down to the coast at Santa Cruz del Sur.

To the east wave energy again declines, this time into the Golfo de Guacanayabo, north of the Manzanillo Peninsula. The delta of the Rio Cauto is growing on the eastern shore of this gulf, north of Manzanillo, and has shores that are partly sandy and partly muddy, with mangroves. The peninsula runs out to Punta Casimba, from which low hilly ridges grow eastward into the Sierra Maestra Mountains, which rise to 1,974 m and descend abruptly to the coast, trenched by short, steep valleys. Terraces indicate intermittent uplift on this bold coast up to more than 300 m above sea level. Santiago de Cuba (> Fig. 5.1.5) stands beside an estuarine inlet as the ranges decline eastward. To the east several ­rivers converge in Guantanamo Bay, a broad ria with the delta of the Guantanamo River on the western shore. The southern part of the bay is occupied by a United States naval station, extending out to Punta Balovento. The coast steepens again eastward to Punta Caleta and the ridges decline to low-lying Cabo Maisi at the eastern end of the island. The northeast coast of Cuba is also steep, backed by the Sierra Nipe-Sagua-Baracoa, with Baracoa at the mouth of the Toa valley. There are nearshore coral reefs to the northwest. To the west are lowlands around a series of swamp-fringed bays, Bahía de Tánamo, Bahia de Levisa and Bahia de Nipe, formed by marine submergence of

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⊡⊡ Fig. 5.1.4 Sandy beach at Ancon.

⊡⊡ Fig. 5.1.5 Santiago de Cuba. (Courtesy Victor Aluija.)

v­ alley mouths and low areas, and extending to the Bahia de Banes. The coast then curves north to Cabo Lucrecia. Westward are more inlets and bays, including mangrove-fringed Bahia de Puerto Padre, Bahia de Malaguata, Bahia de Manati and Bahia de Nuevitas, which marks the beginning of a submerged lowland corridor behind an elongated peninsula that breaks westward into a chain of

reefs, cays and emerged coral limestone islands – the Archipielago de Camagüey, including Cayo Sabinal, Cayo Guajaba and the long Cayo Romano. There has ­evidently been transverse tilting along the coast. Bahia Mayanabo and Bahia La Gloria, separated by a swampy isthmus, continue westward into a long lagoon, the inner (southern) shore intricate, with mangrove swamps, inlets

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and the Laguna de Leche. Offshore is the Canal Viejo de Bahamas, and beyond it the Grand Bank of the Bahamas. Further west the Bahia de Santa Clara continues the elongated lagoon behind another chain of reefs and cays – the Archipielago de Sabana. At its western end the Bahia de Cardenas is sheltered by a large spit that has grown northeast, the Peninsula de Hicacos, with a recurved termination at Punta de Morlas and the remains of an earlier one at Punta Garda. From here past Matanzas to Havana the north-facing sandy beaches are almost continuous, with some low cliffs and a shallow sea, and there are numerous seaside resorts. Beach erosion has occurred here, particularly during stormy winters.

5.1

References Ducloz C (1963) Etude géomorphologique de la région de Matanzas. Cuba. Archives de Sci Genéve 16:351–402 Fairbridge RW (1975) Cuba. In: Fairbridge RW (ed) Encyclopaedia of world regional geology, Vol 1. Dowden, Hutchinson, & Ross, Stroudsburg, PA, pp 252–255 Juanes JL (1996) The erosion in the beaches of Cuba – alternatives for their control. Thesis in option of doctor’s scientific degree in geographical sciences. Havana Khudoley KM (1967) Principal features of Cuban geology. Bull Am Assoc Petr Geol 51:668–677 Rodríguez R, Córdova E (2006) Erosion on the beaches of the northeastern region of Cuba. Pan-American Institute of Geography and History, México

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5.2 Jamaica John Norrman · Tommy Lindell 

1.  Introduction Jamaica is located on the northern margin of the Caribbean Plate, which dips off the narrow northern Jamaican shelf into the Cayman Trench, reaching depths of 7,000 m. The island is 240 km from east to west and 60–80 km front north to south. A central east to west backbone of volcanic and metamorphic rocks forms a discontinuous mountain ridge surrounded by limestone plateaux, which cover 60% of the island. The highest part of the mountain ridge, the steep Blue Mountains, is found in the east with a peak of 2,292 m. To the west of these mountains the hilly landscape is generally about 850 m above sea level, gradually declining to 550 m in the western part (Lindell 1999). There are two major fault systems, one in an east-west direction along the straight central north coast. Raised reefs are numerous, up to a height of 180 m. The other fault system runs mainly in a NW–SE direction throughout the island (Norrman et al. 1997a, b). Along the coast these faults form step-like platforms and cliffs. Variation in geology, landscape morphology and ­climate are the major factors which generate coastal landforms. There is surprisingly large coastal variability, reflecting these factors and the associated vegetation. A series of large alluvial basins in the southern part are occupied by agricultural lowlands and almost inaccessible wetland morass with various kinds of vegetation. The island coasts have a tropical climate. Kingston has mean monthly temperatures of 24.4°C in January and 27.2°C in July, with an average annual rainfall of 800 mm. There is a major contrast between the broad landscape relief between the Blue Mountains in the east and the hills, plateaux and low plains to the west. This contrast, combined with the easterly trade winds, has a strong influence on the climate, resulting in mountain rain forests in the east (mean annual rainfall up to 5,600 mm), and dry savannah type vegetation in the south and west (down to less than 1,000 mm). Torrential rivers carry large quantities of coarse sediment (sand and gravel), derived from the steep eastern mountains and the high hilly areas down to the eastern

parts of the north and south coasts, and the sand and silt derived from slope erosion and agricultural soil degradation are transported by the sluggish streams that drain the wide basins in the south. The tide range on the Jamaican coast is less than 0.3 m; Port Royal has a mean spring tide range of 0.2 m. Coastal processes are dominated by wind-generated waves, and wave energy is unevenly distributed. Comparatively high waves, generated by the trade winds, frequently reach the E, NE and SE coasts, whereas the W coast is sheltered. Terrestrial processes produce sediment ranging from boulders to silt and clay, the coarser material deposited on beaches, while the fine sediment flows out in suspension by the waves to settle offshore or in very sheltered areas along the coast. The trade winds are rather constant, generating a dominant longshore current from east to west along the northern and southern coasts. In bays, a counter-current may occur. Climatic conditions in the Caribbean during autumn and winter favour the generation of hurricanes, which are unpredictable. The hurricane tracks that affect Jamaica usually approach from SSE–SE, south of Haiti, but some tracks pass over the eastern part of Cuba and then turn west to the north of Jamaica. A hurricane passing within about 65 km from the coast generates nearshore waves with a turbulent flow of at least 2–4 m/s (Woodley 1992). In 13 of the last 120 years, hurricane centres have passed close to Jamaica, with a median interval of 6.5 years. Much destruction is caused on the coast by the winds and waves of hurricanes. Low-lying areas are flooded, vegetation damaged and swept away and beaches  re­­moved, but features are restored by natural processes in subsequent calmer weather. However, hard, branching staghorn corals are a special problem as they are brittle, and are ­broken by storm waves, but these do not severely harm the solid reef foundation. According to Wood­ley  (1992) the branching corals grow quickly, and are restored to full maturity in about one to two decades. Man-made structures are vulnerable and expensive to repair. Coral reefs are the marine ecological environment in tropical waters (Degerfeldt and Hendry 1987). They

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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form a natural shore protection, by absorbing wave energy and producing calcareous beach material such as broken ­corals, crustaceans and calcareous algae. The growth of coral reefs depends on non-turbid water and sea depth distribution. On the northern coast, the small rivers west of Annotto Bay carry little sediment. The shelf is narrow, and between Port Antonio and Montego Bay there is an almost continuous fringing reef (> Fig. 5.2.1). In Jamaica where shore-based tourism is of primary economic importance, a sustainable reef system is crucial. Emerged coral

reefs are found locally, as near Oracabessa on the north coast (> Fig. 5.2.2).

2.  The Jamaican Coast On the south coast there are several rivers with a considerable suspended sediment load. Some sections have a broad, somewhat deeper shelf. These factors make coral reefs discontinuous and patchy. Many of the reefs are in ⊡⊡ Fig. 5.2.1 A continuous fringing reef at the outer margin of a shallow platform and a narrow beach of derived coralline sand are found on parts of the northern coast.

⊡⊡ Fig. 5.2.2 Emerged, weathered coral reef east of Oracabessa.

Jamaica

poor condition because of sewage pollution, suspended sediment and excessive fishing. The distribution of fluvially-supplied sediment produces specific provinces of gravel and brown beach sand, in contrast to white beach sand produced by the coral reef systems. There is a brown sand province around the eastern part of the island from the Palisadoes (Kingston) to Annotto Bay (> Fig. 5.2.3), but the white coralline sand beaches around Morant Point form a major enclave in this province, derived from the coral and algal reefs of Morant Bay. White coralline sand derived from a fringing reef, has supplied the beach in Long Bay (> Fig. 5.2.4) on the northeast coast. Copacabana Beach has coarse sand and pebbles, as has the cove in Fairy Hill Bay, which receives gravel from a deep ravine. The fluvial sediment yield from Yallahs River has been reduced by the extraction of gravel upstream, resulting in shore erosion, while erosion of the alluvial fan at the mouth of the river has yielded boulders and gravel to adjacent beaches (> Fig.  5.2.5). Longshore drifting of  sediment supplied by the Morant River has formed a ­barrier beach. The north coast west of Annotto Bay and the west coast are other major provinces of white coralline sand, which reflect the presence of fringing coral reefs  and  only ephemeral rivers. The south-western and ­southern coasts have beaches with various mixtures of white  and  brown sand from rivers carrying sand and silt  and the less luxuriant coral reefs than on the north coast. Some beaches are backed by dunes, as at Thatch­­ field,  and sand has blown inland from Treasure Beach. ⊡⊡ Fig. 5.2.3 In Annotto Bay sand and gravel have been supplied by the Wag Water River, the mouth of which has been deflected towards the southeast by the growth of the barrier spit.

5. 2

Longshore drift has supplied sand and gravel to spits and tombolos, including the barrier spit that protects Kingston Harbour (> Fig.  5.2.6). Groynes have been inserted in an attempt to retain beaches, particularly at seaside resorts. In Jamaica lowland coasts are occupied by wetlands bordering the sea. There are four large coastal wetland areas called the Great Morasses: on the west coast the Negril Morass, on the south coast the Black River and the Portland Bight Morasses, and on the eastern end of the  island the Cape Morant Morass. On the south coast there are also many smaller wetlands, some lagoons and swamps occupying narrow swales between multiple beach ridges. On the north coast there are fewer, smaller wetlands often, degraded by exploitation. The wetland vegetation varies from thick forest to open swamp, and from fresh water to salt water communities, the latter dominated by mangroves. In the outer parts of the Black River and Negril Morasses the swamps are underlain by sedge and mangrove peat for many metres down to the substratum. Radiocarbon dating of peat from different levels has shown that that the swamp surface grew upward at the same rate as the global Holocene sea level rise (Digerfeldt and Hendry 1987), indicating isostatic stability inl these areas for at least the past 8,000 years. Irregular cliffs cut in the rock formations of the Blue Mountain foothills dominate the north-eastern Portland coast, with beaches between cliffy headlands. Another very irregular cliffed coast is found near Hanover, west of

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Montego Bay. Steep limestone cliffs protrude on several headlands along the northern and south-western coast, and there are prominent cliffs along Portland Ridge and the Hellshire Hills. There are abrasion shore platforms on the Negril cliffed coast (> Fig. 5.2.7). About 100 km SSW of mainland Jamaica, Pedro Bank extends over 2,400 sq. km within the 20 m depth contour.

Pedro Cays in the eastern part of the bank area are a group of four tiny coralline islets. The oldest reefs on the Bank form a peripheral strip along the northern edge and there is a swarm of patch reefs at the western end. These reefs are covered by at least 25 m of water, which suggests that they are drowned and support very little active coral growth. At the eastern end of the Bank a third set is

⊡⊡ Fig. 5.2.4 Long Bay on the northeast coast, has a white sand beach with a high berm, behind a fringing reef exposed to dominant easterly waves. (Courtesy L-U. Bergström.)

⊡⊡ Fig. 5.2.5 A wave-cut cliff on the Yallahs River alluvial fan, with a steep shingle beach and residual boulders.

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­ ourishing. The present day shoals are covered with extenfl sive sheets of carbonate sediment, mainly composed of reef detritus and skeletal debris. The low cays are characterised by scrub and grass vegetation growing on coralline sand and gravel, and there are some small beaches with pebbles and sand. Pedro Bank is the largest and most productive fishing ground in Jamaica. The commercial categories include

⊡⊡ Fig. 5.2.6 The southern coast of the Palisadoes. The barrier spit protects Kingston harbour. (Courtesy L-U. Bergström.)

⊡⊡ Fig. 5.2.7 High terrace above steep cliffs. (Courtesy L-U. Bergström.)

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scale fish, lobster and conch. Parts of the cays are permanently occupied by fishermen, who usually stay for a month and return to the mainland for two weeks. Note: All material in this chapter was collected from our project on Integrated Coastal Planning and Management in Jamaica (1992– 1998), a co-operation between the National Resources Conservancy Authority, Jamaica and Uppsala University, Uppsala, Sweden.

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References Degerfeldt G, Hendry MD (1987) An 8,000 year Holocene sea level record from Jamaica: implications for interpretation of Caribbean reef and coastal history. Coral Reefs 5:165–69 Lindell LT (1999) Coastal zone mapping of Jamaica for planning and management. Proceedings, Pecora 14/Land Satellite Information III, Denver, CO

Norrman JO, Lindell LT, Bergström L-U, Mohlund ö, Nisell J (1997a) Manual for integrated coastal planning and management in Jamaica. CBA Internal Report No. 7 Norrman JO, Lindell LT, Bergström L-U, Mohlund ö, Nisell J (1997b) Coastal Atlas of Jamaica. CBA Internal Report No. 10 Woodley JD (1992) The incidence of hurricanes on the north coast of  Jamaica since 1870: are the classic reef descriptions atypical? Hydrobiologia 247:133–138

5.3  Hispaniola

(Haiti and the Dominican Republic)

  Anja Scheffers · Tony Browne  

1.  Introduction Hispaniola is a mountainous island about 660 km long and up to 260 km wide, with the Dominican Republic in the east and Haiti in the west. It has many features in common with the islands of Cuba and Jamaica to the west and Puerto Rico to the east, the island group forming the Greater Antilles. Four nearly parallel WNW-trending mountain ranges are separated by relatively narrow, alluvial-filled, longitudinal structural depressions (> Fig. 5.3.1). Coastal plains are also found, and are most extensive on the southeast coast of the Dominican Republic. In Haiti, where the mountains frequently come close to the shore, the area of coastal plain is relatively small. This account is based on a review by Alexander (1985). The island is located on the northern boundary of the Caribbean plate, and has been subjected to folding, faulting, uplift, and depression since the early Cretaceous (Bowen 1975). Raised coral reef terraces, all of Quaternary age, are found at a number of places along the coast, indicating that local uplift of up to several hundred metres continued at least well into the Pleistocene. Although geological studies of Hispaniola are relatively numerous there have been few accounts of the coastal geomorphology (Barrett 1962). As in Cuba there are sectors of steep, sometimes cliffed, coastline alternating with sandy beaches (some white and calcareous from reefs, others dark and of volcanic origin), and mangroves are extensive on low energy shores, particularly behind shallow coralline seas that are bordered by reefs and island chains. Much of the coast is fringed by marine Quaternary deposits, including emerged coral limestone and associated calcarenites. There is widespread evidence of terraces indicating phases of higher relative sea levels and tectonic uplift in Quaternary times. Hispaniola has a humid to subhumid tropical climate. The north coast is exposed to northeasterly trade winds that bring heavy rainfall in the summer months, but the

south coast is drier. Wave action is strong on the north coast when the trade winds blow onshore, and moderate on the south coast which receives waves generated across the Caribbean Sea. The island coast is also subject to occasional hurricanes, especially in August, September and October. Tide ranges are small: Jacmel on the south coast of Haiti has a mean spring tide range of 0.6 m. The marine provinces bordering Hispaniola are rich in calcareous biogenic sediment, notably coralline sand and gravel, but beaches on the more exposed parts of the coast have been eroding in recent decades.

2.  The Coast of Haiti This discussion of coastal features will start at the international boundary on the north coast and work anticlockwise around the island, first in Haiti, then the Dominican Republic. From the Dominican border to Cap Haitien the coast is backed by a broad alluvial plain, the landward limit of which is a bevelled rock platform possibly formed by marine erosion. The platform ranges in elevation from 150 m in the east to 75 m in the west. Between Cap Haitien and the Baie de 1’Acul the coast is steeply cliffed on Tertiary sedimentary formations. Coral reefs are generally abundant in sheltered bays and bay mouths. Coastal inlets have formed where rivers incised through an emerged fringing reef during Pleistocene low sea level phases. East of Cap Haitien a 20 km long reef stands a little offshore, enclosing a mangrove forest. West of Baie de 1’Acul is a 10 km section of tombolo coast with long, curving beaches. Offshore is Tortue Island, 37 km long and up to 7 km wide. Its entire north coast is dominated by high sea cliffs and is inaccessible. There are two terraces, best preserved at each end of the island, where the slope is relatively gentle. On the southern side the terraces have been virtually destroyed by marine erosion.

This is a revised version of a chapter by C.S. Alexander in The World’s Coastline (1985:181–185). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 5.3.1 Coastal features of Hispaniola. (Courtesy Geostudies.)

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A remarkable set of well-preserved Quaternary terraces occur around the western end of the northwest ­peninsula. They increase in number and size westward from Port-de-Paix, eventually extending to a limestone plateau 640 m high. Successively higher terraces are veneered with progressively older limestones. There are at least 20 terraces between Bombardopplis and Baie de Henne, and they diminish again eastward to Port-a-Piement. There are at least eight high, prominent terraces leading up to the top of the plateau. Uranium series dating of corals from the crests of the three lowest terraces of this series yielded mean dates of 81,000 ± 3,000 years bp for the Mole terrace (16 m), 108,000 ± 5,000 years bp for the Saint terrace (28 m), and 130,000 ± 6,000 years bp for the Nicolas terrace (52 m). Assuming sea level was 6 m higher than at present during the formation of the Nicholas reef, the local uplift rate is calculated to be 0.35 m/1,000 years (Dodge et al. 1983). On the southeastern side of Cap St. Marc there are four prominent limestone terraces which show evidence of crustal tilting. On the northern side of the promontory, two terraces skirt the shore of St. Marc Bay. Except at the head of the bay, the shore is bordered by an almost continuous cliff 20–30 m high. The alluvial lowlands along the west coast of Haiti have flat, marshy shores with fringing mangroves. One of these lowlands, the Cul-de-Sac, has several coral limestone terraces along its inland margins, deposited when a Quaternary sea extended more deeply into the lowland. Subsequently, the terraces appear to have been tilted. The

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Cul-de-Sac is a rift valley between parallel faults, with Port-au-Prince on its southern side. Gonave Island is about 57 km long and up to 15 km wide. Its shores are almost entirely bounded by coral reefs except for the high cliffed section along the northwestern end. Steep cliffs are also found on the southeastern coast of the island and low mangrove shores lie between the cliffed sections. Ten or more well-preserved limestone terraces occur on the north and west coast. West of Port-au-Prince there are cliffs which give place near Grand-Goave to a zone where rapid alluvial deposition dominates the coast, and much of the shore is lined by mangroves and mudflats. A single, slightly arched terrace occurs from Miragoane to slightly west of Jeremie. This north coast of the Jacmel peninsula is cliffed for long sections, the cliffs ranging from 10–18 m in height, ­interrupted by submerged karstic features. Coral reefs are relatively rare and, along with mangroves, are mainly confined to sheltered locations. From Pointe-Dame Marie round to Tiburon a number of small bays are separated by cliffed headlands. The cliffs are occasionally more than 100 m high. At Grante Pointe there is a 10 km long gravel beach 3–5 m high, behind which is a low gravel terrace. The southwest coast of Haiti coast curves out to Point l’Abacou, beyond which are a number of small bays with islands with some detached coral reefs. Between Aquin and Jacmel there are at least two marine terraces, one of which shows evidence of differential uplift. There is also evidence at Jacmel of a third high sea level stand.

Hispaniola (Haiti and the Dominican Republic)

The south coast of Haiti is mostly bordered by steep cliffs. Between Jacmel and the Dominican Republic boundary the cliffs are occasionally 200 m high as the mountains drop to the sea. Coral reefs are relatively rare and ­mangrove growth is confined to a few protected bays.

3.  The Coast of the Dominican Republic From the Haitian border on the south coast of the island cliffs cut in Tertiary limestone run south-east to Cape Beata. There are three terraces in the vicinity of Cabo Beata. The lowest has an elevation of about 40 m, the second lies between 90 and 250 m, and the third has an elevation of about 400 m. Weyl considered that the terraces indicate the extent of Pleistocene uplift in the area. East of Cape Beata the coast trends northeastward, and passes into cliffs in Tertiary limestone. There are low sectors where Quaternary deposits occupy basins east to Punta Palenque. From Punta Palenque to Cabo Engano the coast is almost continuously cliffs of disintegrating limestone, the cliff height varying from less than a metre to about 18 m ⊡⊡ Fig. 5.3.2 Crumbling cliffs of limestone at Boca de Yuma on the southeastern coast of the Dominican Republic near Cabo Espada. (Courtesy W. Barrett.)

5. 3

(> Fig.  5.3.2). Nearly everywhere the cliff is topped by a low ridge of calcareous debris that may obscure the cliff in low sectors. Where prominent, the cliff can be divided into two general levels, a submerged level at about 4 m and an emerged level or secondary cliff at about 3–4 m (Barrett 1962). In places the mountains come close to the sea. Barrett (1962) studied the marine terraces along the southeast coast of Haiti between Punta Palenque and Cabo Engano. The inner edge of the coastal plain lies at an elevation of about 120 m, but clearly defined emerged coast features are absent above 78–82 m. Below 78–82 m are numerous marine-eroded scarps and benches but none are continuous for long distances. However, the scarps and benches can be grouped into a number of ­composite terraces, the most important of which is between 29 and 32 m, extending from Punta Palenque to the bay east of La Romana. The terrace surfaces do not appear to have been warped or tilted and are thought to be of eustatic origin. A variety of evidence from overdeepened river valleys, sand dune remnants, and submerged marine terraces suggests a lowering of sea level by at least 150 m (Barrett 1962). Uranium series dating of corals from the 3–6 m and 8–9 m terraces indicate the age of the lower to range between 95,000 and 120,000 years bp,

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whereas one sample from the upper gave an age of 190,000–210,000 years bp. From Cabo Engano to Cabo San Rafael the coast is low and is bordered by an intermittent coral reef. Much of the shore consists of a calcareous sand beach backed by a low dune, both of which rest on Cretaceous limestone (Barrett 1962). Marine terraces border the south coast of the Bahia de Semana and near Cabo Samana, and there are coral reefs in the bays. The north coast of the Samaria peninsula is steep and rocky with coral reefs, and towards Cabo Francis Viejo the coast is low and bordered by a fringing reef. The Cape has a series of raised coral reefs up to at least 360 m above sea level. In the broad bight between Cabo Francis Viejo and Puerto Plata coral reefs extend as much as 12 km offshore. There are some marine terraces the vicinity of Cabo Francis Viejo and along the coast to Puerto Plata. Sandy and pebble beaches occur at intervals along the north coast of the Dominican Republic west of Cabo Francis Viejo, where the coast is backed by mountainous terrain. At Punta Mangle the coast curves southward and is low, with an almost uninterrupted mangrove fringe as

far as the Haitian border. There is a 1–2 km wide reef platform, separated from the mangrove belt by a reef lagoon several hundred metres wide. Beach monitoring in the Dominican Republic has shown depletion of beaches by storm waves during Hurricane Hugo in 1989 followed by a phase of accretion, then further erosion during later hurricanes (Cambers and James 1994).

References Alexander CS (1985) Hispaniola. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, Stroudsburg, pp 181–185 Barrett W (1962) Emerged and submerged shorelines of the Dominican Republic. Instituto Pan Americano de Geographie e Historie. Geogr Rev 30:51–77 Bowen C (1975) The geology of Hispaniola. In: Nairn EM, Stehli EG (eds) The Ocean Basins and Margin, 3: the Gulf of Mexico and the Caribbean. Plenum, New York, pp 501–522 Cambers G, James A (1994) Sandy coast monitoring: the Dominican example 1987–1992 UNESCO Reports in Marine Sciences 63 Dodge RE, Fairbanks RG, Benninger LK, Maurrasse E (1983) Pleistocene sea levels from raised coral reefs of Haiti. Science 219:1423–1425

5.4 Puerto Rico

Jack Morelock ∙ Wilson Ramirez ∙ Maritza Barreto

1.  Introduction Puerto Rico is 62  km by 178 km, with an area of 9,104  sq.  km and a 501 km coastline (Morelock 1978). There are more than 150 beaches with a total length of 208 km, or 41% of the coastline. The beaches are relatively short, and separated by headlands into distinct beach systems that are closed or semi-closed units receiving sediment from limited local sources and transmitting little sand to other beaches. The geology of the Island is complex, as is to be expected from its formation at the leading edge of the Caribbean Plate. Puerto Rico occupies part of a platform that includes the Virgin Islands. The basic pattern is a volcanic and plutonic central mountain core trending west to east with thick limestones to the north and south (Kaye 1959). The volcanic core has steep, highly eroded ridges separated by deep valleys. The limestone flanks form a series of ridges and valleys that are related to fault trends; they have karst topography and end in steep cliffs along the coast. The coastal plains are at the mouths of river valleys and below colluvial fans from the mountains; they are fringed along the coast by beaches, dunes, beach rock and dune calcarenites. River discharge, rainfall and vegetation cover are greater on the north coast, which has most of the major rivers. The offshore sediment of the narrow (2–4 km wide) north and northwest insular shelf are dominantly terrigenous. The major protection from wave action is provided by remnants of former dune calcarenite lines. Off the semi-arid south coast, the shelf varies in width from 1 to 10 km and has numerous coral reefs. The southwest and eastern areas of the shelf are relatively wide and are predominantly coral and carbonate sediment areas (Morelock 1978). The west and east coasts of Puerto Rico are cut across the pattern of folding and faulting, and the geomorphology reflects both structural trends and rock types. Rocky limestone and volcanic coasts alternate with alluvial deposits,  beaches and mangrove shorelines. The north coast is formed by Tertiary limestones with extensive karst topography and narrow alluvial plains in front of the mountains, where beaches and dunes have developed, in places ­lithified

to beach rock and dune calcarenites. Large areas of mangrove are present, but they are fronted by beaches, spits and barriers. The south coast has similar rock types except for intrusive igneous rocks at the eastern end, and a large area of alluvial fan deposits. Because of lower rainfall, karst erosion on the limestones in the western half of the south coast is much less than on the north coast. The beaches of Puerto Rico contain sand derived from several major sources (Morelock 1978). Transport of rock material by rivers to the beach has supplied terrigenous sand, the composition of which depends on catchment geology. Outcrops of volcanic rocks, notably basalt, have supplied dark minerals and dark rock fragments; weathered granite has yielded quartz and feldspar sand. Beach sediment derived from the erosion of cliffs and shore outcrops are also influenced by the nature of rock outcrops. Sand and gravel swept in by wave action from the sea floor include calcium carbonate supplied to the beach by the shoreward transport of the shells of marine organisms, as well as material eroded from limestones and dune calcarenites. The composition of particular beaches depends on the proportions of these various source materials: near river mouths there is an increase in the terrigenous content of beach sands, and there is also a shift toward terrigenous beaches where an offshore carbonate source is lacking. Storms and hurricanes are the major agent of geomorphic change on the coastline of Puerto Rico (Morelock et al. 2000; Donnelly 2005). A day of storm activity can shape features that would take years to form under normal conditions. The distribution of wave regimes varies around the Island, with changes more rapid on exposed sectors of coastline. Although periods of more than 2 m/year ­erosion have occurred locally, most of the coastline shows alternations of erosion and accretion, with a trend toward moderate erosion of unconsolidated deposits (Liegel 1972).

2.  The Coastline of Puerto Rico The northwest coast near Punta Borinquen is a series of rocky karstic limestone cliffs receding as the result of wave erosion. Where there is a narrow coastal plain in front of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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the limestones, sand from the beach is being incorporated into backing dunes. The west coast of Puerto Rico has the terminations of east-west mountain ridges separated by the broad alluvial valleys of the Culebrinas and Guanajibo rivers (Kaye 1959). The ridges form rocky coastline, and the coastlines of the alluvial valleys are fringed by sandy beaches. The insular shelf is very narrow from Aguadilla to Mayaguez. Much of this coastline receives NW waves refracted from swell generated by north Atlantic storm activity. To the south, headlands alternate with beaches, as at Punta Higuero. Here the beach has sands with almost equal parts of carbonate sand from the shelf and quartz and feldspar, dark minerals and dark volcanic rock fragments from the adjacent coast and river sediment. The sand composition is similar south to Mayaguez, although there is an increase in dark minerals and rock fragments southward. The largest estuary on the west coast is Río Anasco. At Mayaguez the submarine canyon of the Río de Yaguez separates the beaches to the south and north; the  composition changes drastically. The south beach is ­composed of igneous rock fragments and dark minerals, mainly serpentine. The absence of carbonate shell material indicates almost no onshore transport of sediment. South of Mayaguez, the shelf is relatively broad and reefs and shoals with depths of less than 10 m extending 20 km offshore. The pattern of ridges ending in rocky headlands and valley-mouth bays with sandy beach continues southward along the coast, but there are also mangrove shores and fringing reefs because of the protection

offered by the broad shelf and irregular shoal reef bathymetry, which results in reduced wave energy. The beaches south of Punta Guanajibo are of carbonate with quartz and some igneous rock fragments. Punta Arenas has a short beach of carbonate sand (>Fig. 5.4.1) near to the Laguna Joyuda on a coastal plain fronting volcanic hills. From Punta Ostiones south, the sands are largely carbonate, as on bay-head Boqueron Beach, with some quartz derived from rivers draining the alluvial plains east of Boqueron. At the western end of the south coast of Puerto Rico is the Cabo Rojo tombolo, where a low sandy isthmus with a wide beach on its western shore runs out to a headland of Cretaceous limestone (>Fig. 5.4.2). On the south coast of Puerto Rico east of Punta Cabo Rojo the coastline has a narrow fringe of mangroves to Punta Montalva, with low tidal flats and salt ponds to the rear. The irregular cliffy coast east of Punta Montalva is cut in limestone, with small pocket beaches of shelly sand and gravel at the base of the cliffs. Karstic weathering is present, but inhibited by the low rainfall. East from Punta Verraco wave erosion has smoothed sectors of low depositional coast, which are beach-fringed with some low cliffs of alluvial sediment. From Ponce eastward, the coast is a low-lying alluvial plain. There are almost continuous beaches, interrupted by short sectors of mangroves, eroding alluvial plain and rock riprap placed to counter severe erosion. The beach sediment is primarily igneous rock fragment and dark mineral. Many of these beaches have gravelly material in the surf zone, as at Manzanilla (>Fig. 5.4.3). ⊡⊡ Fig. 5.4.1 Carbonate sand beaches at Punta Arenas.

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⊡⊡ Fig. 5.4.2 Punta Cabo Rojo is a block of Cretaceous limestone connected to the mainland by a low sandy isthmus to form a tombolo.

⊡⊡ Fig. 5.4.3 The beach at Manzanillo is lag gravel derived from an alluvial fan.

From Punta Las Marias to the east, the coastline is a low-lying wave-trimmed alluvial plain south of the central mountains. There are extensive narrow beaches of sand and gravel at the base of low cliffs. Mangroves occupy the shore in Bahia de Jobos, where wave energy is low behind a wide zone of coral reefs. At Cabo Mala Pascua granodiorite comes to the coast, and there are small valleys with alluvial fans between headlands. The beaches near the mouth of Río Maunabo are rich in quartz sand.

East of Arroyo is a rocky coastline with small pocket beaches. To the east, the beach sands show an increase in the amount of quartz and feldspar, as well as magnetite and other heavy minerals derived from the large southeast granodiorite batholith. From Arroyo to Naguabo rocky headlands alternate with a few small and narrow beaches. The east coast of Puerto Rico is bordered by a shallow shelf with abundant coral and marine organisms which generate carbonate shelf sands. As a consequence, the

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beaches are again a mixture of carbonate sand from the sea floor and of quartz and feldspar with igneous rock material from alongshore and hinterland sources. All the beaches have relatively high quartz content, reflecting drainage from the weathered igneous rocks of the eastern part of Puerto Rico, but there are varying proportions of calcareous sediment from the sea floor. Morphologically, the east coast of Puerto Rico differs from the west coast in that it is less indented and has broader headlands. Large rocky headlands of plutonic and volcanic rock are located between broad alluvial plains. Small pocket beaches are seen at the base of the cliff in some places. Extensive alluvial plains have been formed by  the major river systems at Naguabo, Humacao and Yabucoa, and there are long beaches between Punta Fraile and Morro de Humacao. Bahia Las Cabezas and Las Croabas have narrow carbonate beaches lying behind a fringing reef which supplies most of the beach material. There are igneous rock fragments in the Las Croabas beach, derived from local sources. Fringing reefs line the coast between Fajardo and Cabo San Juan. The north coast of Puerto Rico is exposed to Atlantic swell and storms arriving across a narrow insular shelf, and has higher wave energy than elsewhere in Puerto Rico. From Cabo San Juan (>Fig.  5.4.4) in the northeast, the coast westward to San Juan is generally low-lying, with beaches and barrier spits at the mouth of the Río de Loiza. The coastal area is low and sandy except for occasional

dune calcarenite bluffs, as at Puerto Vacia Talega. The lowland extends inland for up to 7 km. Several large mangrove areas lie several hundred metres inland behind sandy beaches and barriers. The beach sediment is predominantly carbonate sand from the sea floor, except between Punta Uvero and Punta Vacía Talega, where the beach receives a significant input of river sediment. Quartz content is especially high because the upper drainage basin of the Río Loiza is an area of weathered igneous rock (Gillou and Glass 1957). At the port of San Juan the beaches are thin coverings of sand over a rocky shore. West of San Juan are dune calcarenite cliffs at Punta Maldonado and a large tombolo at Punta Salinas (>Fig. 5.4.5). Much of the north coast is a low-lying coastal plain. There are sandy beaches and dunes with occasionally rocky dune calcarenite coast and beach rock coast west from San Juan. There is a sand spit at the mouth of the Río de la Plata, and to the west is an irregular rocky dune calcarenite coast, which has been breached in many places to form small lunate bays bordered with sand and beach rock (>Fig.  5.4.6). At Porto Nueva a breached headland has been reunited by sand deposition to form a small tombolo. Except for the locally terrigenous sediment supplied by the Bayamon and La Plata rivers, the beach sands are largely carbonate from San Juan to the Río Manatí. West from the mouth of the Río Manatí, the beach sands are igneous rock fragments, dark minerals, quartz,

⊡⊡ Fig. 5.4.4 Cabo San Juan is a headland of volcanic rock with small sand beaches.

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⊡⊡ Fig. 5.4.5 Punta Salinas is a tombolo that has been eroding for the past five decades. Riprap has been placed to counter erosion.

⊡⊡ Fig. 5.4.6 A segment of beach rock at Cibuco indicates the former coastline, and thus the extent of subsequent beach retreat.

and some feldspar. There is a slow decrease in terrigenous material westward until, near the port of Arecibo, the sediment is again carbonate. Sand from these beaches is being delivered to dunes behind the beaches by onshore winds.

The dunes near Arecibo are well vegetated. The beach was sandy into the surf zone until 1999, when storm waves exposed beach rock. Frequent winter storms overwash the dunes carrying gravel on to them. There are karstic cliffs

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west of Camuy, and the dunes west of San Jacinto have been reduced or removed by sand quarrying. Cliffs of Tertiary limestone extend along to the headland at Punta Borinquen.

3.  The Islands of Puerto Rico More than 45 small islands and islets are scattered around Puerto Rico. Some like Isla Magueyes are less than 50 m from the island of Puerto Rico others are 10 (Caja de Muertos) to 100 km (Mona) offshore. The largest three of these are discussed in this section. Vieques and Culebra are populated, Mona is under the supervision of the Department of Natural and Environmental Resources and open to visitors. Isla de Mona, to the west of Puerto Rico (18° N, 68° W) is 7 km by 11 km, and Isla Monito, about 5 km northwest of Mona, is comparatively tiny, only 40 acres in area. Mona and Monito are different from the rest of Puerto Rico. They appear as great flat white slabs of limestone floating on the sea and bounded by high cliffs rising vertically out of the water. They rest on two upthrown structural blocks delineated by several near-vertical faults. On the west coast of Mona are limestone cliffs and fringing reefs. Only on the more protected sides of these islands is the sea sufficiently shallow for the development of coral reefs. On the southwest coast of Isla de Mona they protect a narrow coastal plain partially fringed with sandy beaches. The coastal plain narrows eastward to limestone cliffs, interrupted by low plains east and west of Pajaros, where there is a beach ridge plain and a Pleistocene marine terrace. Reef terraces of late Pleistocene age form the coastline along 14 km of the south and west coasts of Isla de Mona. Beach sands and beach rock extend from Playa Sardinera SE to Playa del Uvero, and from Punta los Ingleses NE to Playa de Pajaros, and slightly beyond. The rest of the coastline is dominated by limestone cliffs. There are caves high in the cliffs at the contact between the grey Lirio Limestone and the underlying buff-colored Isla de Mona dolomite.

The cliffs continue underwater for more than 25 m and shallow sea caves have formed at water level. Isla Vieques (18° 10' N, 65° 20' W) is 13 km off the east coast of Puerto Rico, and measures 5 km by 29 km. The island has Tertiary limestones, Cretaceous volcanics and plutonic intrusive rocks with alluvial plains, beaches and mangrove swamps. Green Beach at the western end is located on a sandy point from which a large shoal submarine sand body extends NW. On the southwest coast, rocky headlands separate pocket beaches and on the south coast curving beaches front mangrove-fringed lagoons. The north coast is rocky, with bordering beaches, some of which are partially protected by fringing reefs. West of Isabel Segunda, rocky headlands separate pocket beaches on the narrow coastal plain. The island of Culebra (18° 20' N, 65° 20' W) is a dissected, partly submerged volcanic island with rocky and beach-fringed coasts and fringing coral reefs. Dakity has a fringing reef on an island at the entrance to Ensenada Honda, a mangrove and rock-lined lagoon on the south coast of Culebra. The north coast is rocky with a few beaches, but on the northern peninsula Flamingo Beach is a striking beach of fine white carbonate sand.

References Donnelly JP (2005) Evidence of past intense tropical cyclones from backbarrier salt pond sediments: a case study from Isla de Culebrita, Puerto Rico. J Coastal Res (Special Issue) 142:201–210 Gillou RB, Glass JJ (1957) A reconnaissance study of beach sands of Puerto Rico. Bull US Geol Surv 1052–1:273–303 Kaye CA (1959) Shoreline features and quaternary shoreline changes, Puerto Rico. US Geological Survey Professional Paper 317-B Liegel LH (1972) NCERI beach-offshore study of mainland Puerto Rico. Puerto Rico Area of Natural Resources, Department of Public Works Morelock J (1978) Shoreline of Puerto Rico. Puerto Rico Coastal Zone Management Program, Department of Natural Resources Morelock J, Capella J, Garcia JR, Barreto M (2000) Puerto Rico – Seas at the millennium. In: Sheppard CRC (ed) Seas at the millennium. Oxford University Press, London

5.4.1 Virgin Islands

Maurice Schwartz

1.  Introduction The Virgin Islands lie at the eastern end of the Greater Antilles. Some, including St. Thomas, St. John and St. Croix, belong to the United States; others, such as Tortola, Virgin Gorda and Anegada, are British. There are high islands with bay beaches and coral reefs rising from the Virgin Island Bank, a submarine platform that runs out eastward from Puerto Rico. St. Croix, an outlier to the south, is located on a submarine ridge south of the Anegada Trough (Mattson et al. 1990). There are also low coralline islands and reefs. Details of one island, St Thomas, are given here.

2.  St. Thomas St. Thomas has an area of 83 sq. km. Rock formations ranging from Cretaceous to Eocene include submarine lava flows, andesite breccias and tuffs, greywackes with blocks of limestone and volcanic rocks, and lenses of foraminiferal limestone. A large batholith intruded during the mid- to late-Eocene was accompanied by tectonic deformation and metamorphism (Weaver 1975; Mattson et al. 1990). There has been faulting and northward tilting, followed by dissection resulting in high ridges with steep slopes, trending mainly NW–SE. There is a wet season from May to November, with NE trade winds prevalent in June and July. This is generally followed by a dry season during the rest of the year, with SE trade winds prevalent from December through March. The trade winds produce high wave energy on eastern coasts, while western coasts remain sheltered. Occasionally, easterly swell generated by Atlantic winter storms reaches the shores of the island. Tropical low-pressure systems form off the west coast of Africa and move westward across the Atlantic Ocean to become tropical storms and hurricanes as they approach the Caribbean Islands. The hurricane season lasts from early June to late November, most occurring in September. The North Equatorial Current coming westward from the Atlantic gives water temperatures of at least 25°C. Mean spring tides are small (0.3–0.5 m) and semidiurnal.

With 85 km of coastline, fringing and barrier reefs on St. Thomas are located predominantly along the northern and eastern shores of the island. A mid-shelf reef complex lies to the south. There are over 40 species of scleractinian corals and 3 species of Millepora in the region; with welldeveloped reefs dominated by Montastraea annularis oc­­ curring at mid-shelf depths of 33–47 m. Surveys have shown that coral reef growth around St. Thomas improves progressing from the inner fringing reefs to the outer mid-shelf reefs. Complicit in this trend are runoff from the land bringing increased pollution and sediment to the shore regions, the Charlotte Amalie harbour being the most impacted of all (Hubbard et al. 2008; Rogers et al. 2008). Added to this around the island is the damage to reefs done by dragging of  anchors and grounding of vessels, careless snorkelers and scuba divers, and overfishing. On a larger scale, hurricanes  and higher sea water temperatures are an added stress  to living coral. Bleaching and disease (black band, white  band, white plague, white pox) also take a toll (Drayton et al. 2005).

3.  Beaches St. Thomas has a number of fine beaches. The sequence described here starts on the southwest coast of St. Thomas and progresses in a counterclockwise fashion around the east end, along the north coast and then ends up at the western tip of the island. Brewers Bay, located near the University of the Virgin Islands, has clear, calm water along a lengthy sand beach. Emerald Beach fronting Airport Road is also known as the Lindberg Bay Beach. Though broad and sandy, in stormy times the beach face is cut by low scarps. To the east, across Charlotte Amalie, the Morningstar Bay beach has fine sand at the upper foreshore but coarse sand and pebbles in the sloping surf zone. Limetree Beach is located in a cove within Frenchman Bay. The sandy beach gives way to a rocky substrate in shallow water. A small stream crosses the beach at one end. While the Bolongo Bay beach is composed mostly of sand and pebbles, beach rock outcrops in places on the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.4.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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foreshore. Along one side of the bay, the shore is rocky. Nestled between Breverhout Point and an adjoining headland, Secret Harbour has a sandy beach and a calm wave climate. The shoreline is fairly close to the base of a restaurant building and palm trees, hinting at some beach erosion over the years in this protected embayment. The beach at Cowpet Bay is sandy with some pebbles and coral fragments. While the backshore is lined with palm trees, rocky areas bracket the ends of the beach. The water is known to be calm and clear. Bluebeards Beach and Turtle Cove are sandy with calm water just offshore. Sea grass is common at both sites. Vessup Beach is at the SW end of a long embayment. Ranging from narrow to wide, the sandy beach has outcrops of beach rock exposed on the lower foreshore (>Fig. 5.4.1.1). Sapphire Beach is sandy with some pebbles, and driftwood occasionally accumulates at the backshore. Shallow reefs on one side and an artificial pond on the other bracket the beach. Lindquist Beach is also sandy, with outcrops of beach rock at intervals (>Fig.  5.4.1.2). A dense growth of mangroves backs some sections of the beach. The narrow beach at Sugar Bay has trees growing at the  upper swash limit and slight scarping. Patches of ­different  sand sizes along its length indicate possible replenishment. The beach at Water Bay, also known as Pineapple Bay Beach and Renaissance Beach, is composed mostly of sand with some small pebbles. Sea grass is abundant offshore

and often litters the beach. Coki Beach is small, with clear calm surf and a nearshore reef. The Mandahl Bay beach is composed of sand, gravel, and coral fragments; there are small tide pools, and some rocky shore areas. Two rock jetties have been constructed at the east end of Mandahl Beach beside a channel into the lagoon. The most noted beach on St. Thomas is that at Magens Bay, deeply embayed between the long, NW–SE trending, wooded headlands of Tropaco Point on the west and Peterborg Point on the east. Magens Beach is a 1.5 km long pocket beach, with fine sand, a gently sloping nearshore bottom: there is no longshore current. Somewhat to the west of Magens Bay there is a small beach at Hull Bay. There is a boat ramp crossing the foreshore at the center of the beach, where sediment has accumulated on the east side and slight erosion has occurred on the west. The beach at Dorothea Bay, embayed between two small, rocky headlands, has a long northeasterly fetch that results in strong wind and wave action (>Fig. 5.4.1.3). The beach sediment mixture of sand and gravel is often formed into large beach cusps in the pounding surf. Neltjeberg Beach a long, broad, expanse of sand with gentle surf. Extensive reefs lie offshore. The west end of St. Thomas has beaches in two coves separated by a promontory, Botany Bay and Sandy Bay respectively. These sandy beaches at the western end of St. Thomas are sheltered from the prevailing northeasterly and southeasterly trade winds, providing a gentle wave environment.

⊡⊡ Fig. 5.4.1.1 Beach rock on the lower foreshore at Vessup Beach.

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⊡⊡ Fig. 5.4.1.2 Lindquist beach.

⊡⊡ Fig. 5.4.1.3 Beach cusps at Dorothea Bay beach.

References Drayton N, Rogers C, Devine B (2005) The state of coral reefs of the U.S. Virgin Islands. The Ocean Conservancy, Washington DC, 58 p Hubbard DK, Burke RB, Gill IP, Ramirez WR, Sherman C (2008) Coralreef Geology: Puerto Rico and the US Virgin Islands. In: Riegl, Dodge RE (eds) Coral reefs of the USA: coral reefs of the world, Vol 1. Springer, Berlin, pp 263–302 Mattson P, Draper G, Lewis JF (1990) Puerto Rico and the Virgin Islands. In: Dengo G, Case JE (eds) The Caribbean region: the geology of

North America, Vol H. Geological Society of America, Boulder, CO, pp 112–120 Rogers CS et al (2008) Ecology of coral reefs in the US Virgin islands. In Riegl BM, Dodge RE (eds) Coral reefs of the USA: coral reefs of the world, Vol 1. Springer, Berlin, p 303–373 Weaver JD (1975) Virgin Islands. In: Fairbridge RW (ed) Encyclopedia of world regional geology part 1: western hemisphere. Dowden, Hutch­ inson and Ross, Stroudsburg, pp 654–656

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5.5  Lesser Antilles

Gillian Cambers

1.  Introduction The Lesser Antilles represent an arc of island systems separating the Caribbean Sea from the North Atlantic Ocean. They extend from the Virgin Islands (latitude 18° 30' N) in the north to Trinidad (latitude 10°N) in the south (Deane 1985). The arc of islands lies along the eastern boundary of the Caribbean Tectonic Plate, which lies adjacent to the North American Plate, which in turn is moving westwards and subducting beneath the Caribbean Plate. The Lesser Antilles consist of two main arcs: an inner active volcanic arc and an outer calcareous arc. The inner volcanic arc is of Eocene origin and extends from Saba in the north to Grenada in the south. Many of these islands have experienced recent volcanic eruptions, as at Mount Pelée in Martinique in 1902, Soufrière in St. Vincent in 1979, and Soufrière Hills in Montserrat in 1995 where the volcanic activity is continuing into the 21st century (>Fig.  5.5.1). The submarine volcano, Kick ‘em Jenny, 8 km north of Grenada, rose from a depth of 200 m

to 140 m below sea level between 1974 and 1989 and may emerge as the next Caribbean island (Hendry 1996). The outer arc, which has been in existence since the early Tertiary, stretches from Sombrero past Anguilla to Barbados in the south. The islands of this arc are much lower and have less rugged topography than those of the inner volcanic arc. Coral reefs have developed around many of these islands and the extent and type of reef is influenced by aspect, substrate and hurricane influence (Wells 1988). Reef development is usually more extensive on the small low islands with low rainfall and little sedimentary runoff, as in Anegada, Antigua, Barbuda and the southern Grenadines. St. Vincent and the Grenadines and the US and British Virgin Islands are archipelagos, and many of the other territories consist of several islands and cays, e.g., Anguilla. The > Table 5.5.1 lists the island territories, located in (>Fig. 5.5.2). The climate is tropical, with waves generated by the Northeast Trade Winds, which are strongest from June to July and from December to March. Waves generally

⊡⊡ Table 5.5.1 Island Territories of the Lesser Antilles Political grouping

Outer chain (Coralline)

Inner chain (Volcanic)

Independent Island States

Antigua and Barbuda Barbados

Dominica Grenada St. Kitts and Nevis St. Lucia St. Vincent and the Grenadines

Departments of France

St. Martin

Guadeloupe Martinique St. Barthélemy

Netherlands Antilles

St. Maarten

Saba St. Eustatius

UK Dependent Territories

Anguilla

Montserrat

UK Overseas Territories

Other Trinidad and Tobago

British Virgin Islands U.S. Virgin Islands

A revised version of a chapter by Compton Deane in The World’s Coastline (1985):193–199. With contributions by Maurice Schwartz (St. Martin) and Pierre Saffache et al. (Guadeloupe). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_5.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 5.5.1 Since the 1995 eruptions of the Soufrière Hills volcano, the southern part of Montserrat, including the former capital Plymouth, has been abandoned. (Courtesy Angela Greenaway.)

approach from directions between northeast and southeast. The east facing coasts are known as windward coasts and experience higher wave energy than the west-facing leeward coasts, which are more sheltered. Between the months of October and April, the Lesser Antilles are periodically affected by swell waves generated by intense midlatitude storms in the North Atlantic Ocean. During each swell season, there may be five to ten swell events each lasting one to eight days and these bring high swell waves to east-, north- and west-facing coasts. One such event in March 2008 resulted in swells more than 10 m in height and caused at least one drowning death in the Lesser Antilles. Between June and November, the Lesser Antilles are affected by tropical storms and hurricanes. In general, these low pressure systems form towards east of the island arc and travel in a westerly or northwesterly direction, although there have been exceptions: Hurricane Lenny in November 1999 formed off the coast of Nicaragua and moved eastwards across the Caribbean Sea to impact the Lesser Antilles. These low-pressure systems bring intense rainfall, high winds and storm surges to the islands of the Lesser Antilles causing significant coastal and beach changes, environmental damage, flooding, loss of life and damage to island infrastructure. During the last century, there have been multi-decadal fluctuations in hurricane

activity in the Atlantic Basin and Caribbean Sea, with the period 1970–1994 showing low hurricane activity, and the period since 1995 showing increased activity. The record is insufficient to indicate whether these fluctuations are a natural climate variation or linked to climate change (IPCC 2007). The period 1995–1999 saw intense hurricane activity in the northern islands of the Lesser Antilles, while hurricane activity moved to the southern islands (Grenada and St. Vincent and the Grenadines) in the years 2004–2005. Tides are semidiurnal and the range is small, around 0.3 m and 0.6 m at Spring tides. Coastal landforms include cliffs, rocky slopes, dunes, beaches, wetlands and salt ponds, beach rock ledges; offshore there are coral reefs, seagrass beds and sandy bottoms. Beach composition varies from large granite boulders to giant coral heads, and from black volcanic sand to white coralline sand. Most of the ports and tourism infrastructure are located close to the west coasts of the islands in the Lesser Antilles. These are the more sheltered coasts and favoured for tourism development. In many places, this infrastructure interrupts coastal processes and results in localised erosion and coastal changes. Selected islands are considered in more detail in this chapter to illustrate the range and variety of coastal geomorphology in the Lesser Antilles.

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⊡⊡ Fig. 5.5.2 Map of the Lesser Antilles. (Courtesy Caribbean Conservation Association.)

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2.  Anguilla Anguilla, one of the most northerly of the outer arc islands, lies on the extensive Anguilla Bank. The island is low-lying and composed of a Lower Miocene limestone and marl series overlying volcanic rocks (Christman 1953). There are several small uninhabited cays. There are no surface streams but several salt ponds resulting from the closure of former bays. The beaches for the most part consist of calcareous sand and shells from the reefs. The island trends northeast/southwest and the predominant littoral drift is from east to west, although several of the bays form partially closed beach cells. Stretches of the north coast are characterised by high cliffs and there are also pronounced embayments. At Shoal Bay, there is a depositional promontory that is eroding, especially at the northern point. Along the north coast, west of Road Point, beaches have formed between limestone outcrops at Road Bay, Long Bay, Meads Bay and Barnes Bays, those at Road Bay and Meads Bay being barriers that have enclosed salt ponds. Along the western part of the south coast, beaches have formed barriers that enclosed salt ponds at Ren­dezvous, Cove, Maundays and Shoal Bays. Extensive dunes, up to 10 m high, developed behind these sandy barrier beaches. However, many of the dunes have disappeared because of alteration for tourism development, past mining for construction sand, and damage in recent hurricanes, particularly between 1995 and 1999. As a result, these barrier beaches are now being breached even during minor weather

events, as at Cove Bay. At Maunday’s Bay successive attempts to dredge and renourish the beach after hurricanes in 1995, 1998 and 1999 were largely un­­successful (>Fig. 5.5.3), and a line of tourism resort villas built on top of a dune system were  eventually protected with a sea wall (>Fig.  5.5.4). Accretionary features such as a tombolo west of Blowing Point Ferry Terminal on the south coast have all but disappeared in recent years. Monitoring of beach profiles over the period 1991–2000 showed an average erosion rate in excess of 1 m/yr (UNESCO 2003). During this period Anguilla was impacted by three hurricanes, which have damaged its extensive coral reefs; after Hurricane Luis in 1995 surveys showed large expanses of Elkhorn (Acropora palmata) reefs reduced to coral rubble (Bythell et al. 1996).

3. St. Martin

(Contributed by M.L. Schwartz) Department of Geology, Western Washington University, Bellingham, Washington, USA

St. Martin, south of Anguilla at the northern end of the  Lesser Antilles, has an area of 95 km2. The world’s smallest bi-national island, it is divided into Saint Martin (France) in the north (capital Marigot) and Sint Maarten (Netherlands Antilles) in the south (capital Philipsburg). The two divisions are known locally as the Dutch Side and the French Side. The island stands in the west-central part of the Anguilla Bank and forms part of the non-volcanic arc of ⊡⊡ Fig. 5.5.3 Beach nourishment in progress in 1996 at Maunday’s Bay, Anguilla, after Hurricane Luis (1995).

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⊡⊡ Fig. 5.5.4 Sea wall under construction in 2000 at Maunday’s Bay, Anguilla, after Hurricane Lenny (1999).

the Lesser Antilles, which curves east of the volcanic inner arc (Brouwer 1975). It has a basement of silicified tuffs with coarse pyroclastic material including Eocene carbonate fragments, intruded by Oligocene granodiorite. The hard tuffs persist as mountain areas, bordered by hilly country on less resistant igneous rocks, and there are intrusive dykes, stocks and plugs (Christman 1953). An unconformity separates the Upper Oligocene to Upper Miocene limestones, which form the lowlands. Similar to Anguilla the coastline is embayed, but has  been shortened by the formation of barrier beaches enclosing lagoons. There is a double tombolo impounding Simpson Bay Lagoon in the southwest. Pleistocene coral reef terraces occur 5–6 m above sea level, and there are nearshore reefs. Limestone headlands separate sandy beaches. The Dutch Side begins south of the western Terres Basses Peninsula on the isthmus south of the Simpson Bay Lagoon. Cupecoy has low (3–6 m) limestone cliffs facing southwest with narrow beach segments (>Fig.  5.5.5), where a formerly wide beach was depleted by Hurricane Luis in 1995 and Hurricane Jose in 1999. Mullet Bay to the east has low karstic limestone shores. At Maho an emerged limestone terrace is fronted by cliffs with caves, declining eastward to a curving sandy beach with segments of grassy backshore terrace. South-facing Simpson Bay is backed by a 1.5 km long curving sandy barrier, with a bridge across an inlet to the Simpson Bay Lagoon. Promontories and bays continue eastward to Great Bay, where Philipsburg is built on a barrier enclosing a lagoon, the Great Salt Pond, where salt was extracted formerly. In 1995 and 1999 Hurricanes Luis and Lenny damaged beaches and coral

reefs, and the Great Bay Beach was subsequently replenished with dredged sand. Pointe Blanche is a wide limestone peninsula on the southeast coast. The east coast, more exposed to the Northeast Trade winds and strong waves, has limestone cliffs and sandy beaches. A sand spit almost encloses a small lagoon at Oyster Pond. North of Oyster Pond, the boundary between the Dutch and French sides crosses the coast. At Orleans a sandy barrier encloses another lagoon, and on its seaward side is Galion Beach, often strewn with seagrass. Orient Bay, to the north, has a broad sandy beach. Offshore are extensive coral reefs in a marine reserve. On the northwest coast, Grand Case has another barrier beach enclosing a lagoon, and in Friars Bay to the south there are beach gravels (shingle) beneath the sand. South of the town of Marigot the barrier beach in Baie Nettle, on the northern side of Simpson Bay Lagoon, has been extensively armoured against erosion. To the west is the Terres Basses Peninsula, with Baie Longue on its southern coast. This is a wide sandy beach backed by a small lagoon and fronted by a shallow bay where wave energy is generally low.

4. Guadeloupe

( Contributed by Dr. Pascal Saffache and Garry Ramedine) Université des Antilles et de la Guyane

The Guadeloupe Archipelago has 680 km of coastline and is composed of eight islands: Basse-Terre, Grande-Terre,

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⊡⊡ Fig. 5.5.5 Limestone cliffs at Cupecoy Beach, St Martin.

Terre-de-Haut, Terre-de-Bas, Marie-Galante, La Desirade, Saint-Martin and Saint-Barts. While some of the islands are exclusively volcanic (Basse-Terre and the Saintes ­archipelago) in others the volcanic bedrock is covered with calcareous sediment (La Desirade, Grande-Terre, Saint-Martin, Saint-Barths, Marie-Galante). Emergence, due to tectonic uplift, results in cliffs over 30 m high on eastern coasts with sandy coves containing emerged beach rock. Beach rock also outcrops extensively in many other islands of the Lesser Antilles indicating the former position of beaches as coastlines have retreated inland. The Gates of Hell in the north-east is a limestone calanque bordered by flat ledges and backed by a sandy beach washed by Atlantic waves. The western coasts have low cliffs cut in pyroclastic tuff backing beaches and swamps. A tidal channel, Rivière Salée, runs across the northern end of Grande-Terre from the Atlantic to the Caribbean. There are alternating black volcanic sediment and white coralline sediment beaches. Bays and lagoons, e.g., the Grand-Cul-de-Sac Marin and Petit-Cul-de-Sac Marin, the Saint-Martin and Saint-Barths lagoons, are occupied by mangrove swamps. The inner wetlands area is composed of red mangrove (Rhizophora mangle), the middle section by black mangroves (Avicennia germinans), and the outer portion by red mangroves. Sandy beaches in the Guadeloupe archipelago occupy 19% of the coastline, mainly in bays or behind reefs. Most of the sandy beaches are along the south coast of Grand-Terre (between Gosier and Pointe des Chateaux),

in the north-west of Basse-Terre (between Bouillante and Deshaie) and on the south-west coast of Marie-Galante and Desirade. Gravel beaches with black pebbles of volcanic origin are present in bays along the southern coast of Basse-Terre, while white pebbles, formed from calcareous rock, are found only at Anse Pistolette in the north of Grande-Terre. In nearshore water up to 30 m deep, there are fring­ ing  reefs, barrier reefs (such as the Grand-Cul-de-Sac Marin) and volcanic rock outcrops covered with coral. In Guadeloupe, the fringing reef is limited to 90 km in length. The reefs are located mostly along the southern coast of Grande-Terre, in the Petit-Cul-de-Sac Marin, to the south of Marie-Galante, and on the Atlantic coast between Pointe des Chateaux and Port-Louis as well as surrounding Petite Terre and La Desirade. The Grand-Cul-De-Sac Marin barrier reef is about 30 km long and up to 10 km wide, enclosing a 78 km2 lagoon with a maximum depth of 20 m. Behind the solid coral reef is a zone of seagrass (Thalassia testudinum) on sandy shoals, then mangroves along the coast. All these features have been modified by human activities.

5.  Dominica Dominica, with an area of 751 km2 is a rugged, steep volcanic island with a mountainous interior and many rivers. The island began to form during the Miocene period and the volcanic rocks are mainly andesites with subordinate

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dacites and basalts. Active volcanism is evident in the Sulphur Springs region and the area of the Boiling Lake and the Valley of Desolation. Most of the flat or gently sloping land exists near the coasts and this is where most of the urban and agricultural developments have taken place. Most of the 153 km of coastline is cliffed with beaches along 24 km. With a very narrow coastal shelf, there are few true coral reefs in Dominica. Along the west coast, there are a number of narrow beaches made up of a mixture of volcanic sand and stones. Many of the beaches are located near the place where rivers discharge into the sea. They often exhibit a seasonal change in composition with sand during the lower wave energy months from May to September, and stones during the high wave energy months of November through to March. At the SW tip of the island the small islet of Scotts Head is joined to the mainland by a narrow gravelly tombolo (>Fig. 5.5.6). Prior to Hurricane David in 1979, this tombolo was composed of sand and pebbles; it was wide enough to have a football pitch and there were many large trees. Now the tombolo is low and narrow and often breached during storms and hurricanes. Along much of this west coast, the coastal highway runs near to the sea and had to be protected with extensive seawalls, usually resulting in the loss of any narrow beaches that may have existed in the past. This has created problems for land crabs that annually migrate to the sea to deposit their eggs. Monitoring of beach profiles along the coasts of Dominica has shown erosion rates of 1 m/yr over the period 1987–2000 (UNESCO 2003). The capital city of Roseau has been protected with a massive sea wall combined with boulder protection. ⊡⊡ Fig. 5.5.6 Tombolo joining Scotts Head Islet to the main island of Dominica.

5.5

At the northern end of the west coast there is a ­ ronounced embayment, Prince Rupert Bay, where the p second largest town, Portsmouth is located. A once wide, continuous sandy beach now exists only in patches. Behind the beach there is the extensive Indian River wetland area. Further to the north, in the Cabrits Peninsula, once sandy beaches at Belle Hall and Toucarie, which in the past were favoured for recreation by Dominicans, are now eroded and reduced to stones and gravel following several hurricanes between 1989 and 1999. The north coast from Hampstead to Pagua has several sandy embayments and in places such as Pointe Baptiste there are very scenic cliffs and pockets of brown sand beaches. Much of the east coast is cliffed with long stretches of black sand beaches at Londonderry, Castle Bruce and Bout Sable, to mention a few. Most of this coast experiences high wave energy for much of the year. At Marigot a man-made fishing harbour has been created.

6.  Barbados Barbados has an area of 430 km2 and a coastline 92 km long. It consists mainly of Tertiary limestone, marl and flysch deposits, with a Pleistocene coral limestone cap up to 100 m thick. The island has a regular coastline with no deep indentations or offshore islands. The coast consists mainly of long stretches of sandy beach, except in the north and southeast where cliffs rise straight from the sea. On the east coast the Pleistocene coral limestone that covers the rest of the island has been eroded away ­exposing

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the less resistant underlying strata, mainly marls and clays, in a region known as the Scotland District. A 10 km long, wide quartz sand beach backed by low sand dunes occupies most of this coast, which is fully exposed to the Atlantic waves and has minimal protective coral reefs. Blocks of limestone from the eroded cap are visible in the offshore area. This beach ends in cliffs at both ends. At the southern end, the coast becomes more indented with bays separated by cliffs. At Bathsheba there are several stacks (>Fig. 5.5.7) and between Bath and Consett Bay, remains of the old railway line can be seen near low water mark, indicating the erosion taking place along this coast. The SE coast from the Ragged Point lighthouse to South Point is characterised by coral limestone cliffs with several significant bayhead beaches, of which Foul Bay and Long Bay are the longest. The beaches are made up of sand and coral, shell, and algal fragments. The Cobblers Reef, a barrier reef, lies between 1 and 3 km offshore, and encloses a partially sheltered lagoon. Along the south coast from South Point to Bridgetown the coastline is low and rocky with intermittent coralline sand beaches. Offshore there are rubble banks and scattered coral patch reefs. At Oistins a long jetty originally planned for use by the Coastguard trapped a wide sandy beach, known locally as “Miami Beach”. In several places, groynes have been built to trap the westward moving sediment and at Accra Beach there is a submerged offshore breakwater. Near St. Lawrence Gap, the one remaining

significant wetland in Barbados, Graeme Hall Swamp, is connected to the sea via a sluice gate. The coastline is densely developed with tourism properties and associated infrastructure. At the western end of the south coast, a former accretionary feature has been stabilised with a breakwater at Needhams Point. The main town, Bridgetown and the deep water harbour are located at the SW end of the island, and significant accretion has taken place north of the deep water harbour at Pelican Beach. The west coast consists of a low sand terrace varying in width from a few hundred metres to 1 km. There is a nearly continuous coralline sand beach, consisting of some 19 beach cells that are anchored by fringing reefs at the sandy headlands separating the cells (Bird et al. 1979). These fringing reefs are poorly developed and have been stressed for decades by runoff and pollution from densely developed west coast tourism. Some of the beach cells are well developed with wide beaches while others have only narrow strips of sand, which disappear at high tide. Many beachfront property owners have resorted to the use of backshore protection, with varying degrees of success. Despite these pressures, the west and south coasts are important sea turtle nesting sites, particularly for Hawksbill turtles (Eretmochelys imbricata). The beaches of Barbados have been under pressure from erosion (>Fig.  5.5.8) and intensive use since the 1980s and as a result, a Coastal Zone Management Unit has been established to monitor and manage these issues. ⊡⊡ Fig. 5.5.7 Stacks at Bathsheba, Barbados.

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⊡⊡ Fig. 5.5.8 Eroded west coast beach at Heywoods, Barbados.

7.  St. Vincent and the Grenadines St. Vincent is one of the youngest of the islands of the inner volcanic arc. The northern end of the island is formed by the active volcano, Soufrière, which is a recent volcano built up within the last half million years, and was last active in 1979. This volcano, like many of the others in the Lesser Antilles, is monitored by the Seismic Research Unit of the University of the West Indies in Trinidad and Tobago. The island has a central chain of mountains, which are the eroded remnants of a series of volcanoes, and a coastal plain of varying width. Cliffs have been cut in the volcanic formations, and locally there are outlying stacks (>Fig. 5.5.9). The chain of Grenadine islands to the south of mainland St. Vincent has a more complex geological history and has rock types of both sedimentary and volcanic origin (Caribbean Conservation Association 1991). Shore shelf around the St. Vincent mainland is very narrow except at the southeast side. Most of St. Vincent’s beaches are made up of black sand, stones and boulders; the exception is on the south coast where there are coral reefs and white sand beaches. The east coast consists of long stretches of beaches separated by headlands of more resistant rock. The beach and nearshore sediment along the windward (eastern) coast of St. Vincent are dominated by volcanic ash fragments and

heavy minerals (Adams 1968). There are several ­significant rivers, bringing sediment to the coast, e.g. the Rabacca River north of Georgetown, where the river bed is still mined for construction material (>Fig. 5.5.10). Settlements on the coastal plain, near the mouths of these rivers, are vulnerable to flooding during storms and hurricanes (>Fig.  5.5.11). There used to be extensive black sand dunes at Brighton and Diamond Bays at the south end of the east coast. The Brighton dunes, with maximum elevations of 60 m, were formed when sand accumulated north of a tombolo opposite Milligan Cay, and were blown landward by the Trade Winds. Dunes 6 m high at Diamond Bay were removed by quarrying in the 1970s and 1980s. The south coast consists of several beach segments with a mixture of coralline (white) and terrestrial (brown) sand. This coast has been substantially modified with several sea defence structures. The west coast consists of a series of indented bays with volcanic sediment (sand, gravel and boulders). Mining from the beaches and river mouths, although usually small scale, is still a significant issue in St. Vincent. The Grenadine Islands, to the south of mainland St. Vincent, are located on a platform, 40 m deep and are an important yachting destination. There are extensive reef systems, which have been impacted by anthropogenic impacts, hurricanes and coral diseases. Several of the

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⊡⊡ Fig. 5.5.9 South Stack, a volcanic neck with an abrasion notch at present sea level. (Courtesy J.R.V. Prescott.)

⊡⊡ Fig. 5.5.10 River mouth on the east coast of St. Vincent. The river is delivering sand and gravel to beaches, but the yield has been diminished by extraction. (Courtesy J.R.V. Prescott.)

islands have significant wetlands, such as The Lagoon on the SW coast of Mustique. The islands are rocky and vary in size and elevation. The coastlines are indented, with white sand beaches separated by cliffed headlands. There are few significant sand dune features. Bequia, one of the largest islands, has a large embayment on the west coast, Port Elizabeth. This is an important yacht anchorage, and the bay is fringed with a narrow white sand beach.

8.  Trinidad and Tobago Trinidad and Tobago with a combined area of 5128 km2 is situated on the continental shelf off the north coast of Venezuela. Trinidad has three mountain ranges orientated east–west and numerous rivers drain the upland areas ending in coastal swamps and lagoons. Tobago is rugged with an internal mountain range and a narrow coastal plain.

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⊡⊡ Fig. 5.5.11 Villages behind this black sand beach on the east coast of St. Vincent are vulnerable to flooding by seawater and land runoff.

⊡⊡ Fig. 5.5.12 Caroni Swamp, Trinidad, home of the Scarlet Ibis.

The east coast of Trinidad is dominated by three long bays: Matura, Cocos and Mayaro Bays. The beaches at Matura and Mayaro Bays are composed of quartz sand eroded from cliffs cut in Miocene and Pliocene formations and moved by rivers to the coast. At Cacos

Bay, a 20 km barrier beach has formed, impound­‑ ing three rivers to make the large freshwater Nariva Swamp, 70 km2 in extent. The barrier beach is broken only at the centre, where the Nariva River reaches the sea.

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Longshore movement of sand along the south coast of Trinidad ends in the large Icacos spit at the SW tip of the island. The major source of beach sediment is from erosion of cliffs composed of sands, clays and weakly cemented sandstones of Miocene, Pliocene, and Pleistocene age on the western part of the south coast. Mangroves are extensive. On the SW coast of Trinidad cliffs in Miocene sands and clays are being eroded. This sector of coastline comes to an end at Los Gallos, where there are stacks of slightly more resistant Pleistocene sandstone. The near-estuarine conditions of the Gulf of Paria restrict coral development and the coast is dominated by mangroves fringing the extensive Caroni Swamp (>Fig. 5.5.12). Along the north coast, there are several small embayments and considerable coral development. Tobago has a more indented coastline, with several sandy embayments separated by lengths of cliff coastline. The Buccoo Reef area on the SW coast has extensive seagrass and associated mangrove stands.

References Adams RD (1968) Distribution of littoral sediment on the Windward Coast St. Vincent, West Indies. Transactions of the 5th Caribbean Geological Conference, pp 55–59

Bird JB, Richards A, Wong PP (1979) Coastal subsystems of western Barbados, West Indies. Geogr Ann 61A:221–236 Brouwer GC (1975) Netherlands Antilles. In: Fairbridge RW (ed) Encyclopedia of world regional geology, part 1: western hemisphere. Dowden Hutchinson and Ross, Stroudsburg, pp 362–365 Bythell JC, Cambers G, Hendry MD (1996) Impact of Hurricane Luis  on  the coastal and marine resources of Anguilla. Summary report prepared for the UK Dependent Territories Regional Secretariat, 13 p Caribbean Conservation Association (1991) St. Vincent and the Grenadines country environmental profile, 222 p Christman RA (1953) Geology of St. Bartholomew, St. Martin and Anguilla, Lesser Antilles. Geol Soc Am Bull 64:65–96 Deane C (1985) Lesser Antilles: Virgin Islands to Trinidad. In: Bird ECF and Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 193–199 Hendry MD (1996) The geological legacy of small islands at the Caribbean–Atlantic boundary. In: Maul GA (ed) Small islands marine science and sustainable development. American Geophysical Union, Coastal and Estuarine Studies, vol 51, pp 205–224 IPCC (Intergovernmental Panel on Climate Change) (2007) Climate change 2007: the physical science basis. Summary for policy makers UNESCO (2003) Wise practices for coping with beach erosion in the  islands of the eastern Caribbean. 10 booklets, Environment and Development in Coastal Regions and Small Islands. UNESCO, Paris Wells SM (ed) (1988) Coral reefs of the world. Atlantic and Eastern Pacific. UNEP Regional Seas. IUCN Monitoring Centre, 373 p

6.0 Atlantic Ocean Islands – Editorial Introduction

The Atlantic Ocean was initiated about 140 million years ago in the Late Jurassic when the disruption of the ancient megacontinent of Pangaea was followed by westward movement of the American Plate and eastward movement of the Eurasian and African Plates. It is still widening by up to 10 cm/year. The divergence was marked by the formation of the mid-Atlantic Ridge by the extrusion of basalt, resulting in the formation of a series of volcanic  islands extending from Jan Mayen Island in the north  to  Bouvet Island in the south: the series includes Iceland, the Azores, St Peter and Paul Rocks, Ascension, St Helena, Tristan da Cunha and Gough Island. In addition, Greenland is a largely ice-covered island of continental Pre-Cambrian and younger rocks projecting into the NW Atlantic and the Falkland Islands are also of continental rocks, with a Pre-Cambrian basement overlain by Mesozoic formations on the South American shelf. The >Bahamas and >Bermuda are islands of limestone and associated calcareous deposits rising from the American shelf. East of the mid-Atlantic ridge the Faeroe Islands, St Kilda, Madeira, the Selvagens archipelago, the Canary Islands, the Cape Verde Islands and the islands in the Gulf of Guinea all surmount the Eurafrican shelf (>Fig. 6.0.1). The Atlantic islands are generally high islands, mostly of volcanic origin, with cliffed, rocky and boulder-strewn coasts (Hansom 2005). Coral reefs are poorly developed (except in the Bahamas and Bermuda), and the low coralline islands and  cays found in the Indian and Pacific Oceans are not present. There are active volcanoes in Iceland, in the Azores and on Tristan da Cunha. Volcanic structures have influenced coastal evolution, with rapidly receding cliffs in ­scoria and ash while lava flows, dykes and necks form persistent bold cliffs, rocky headlands and stacks. Steep coasts and plunging cliffs are common. Encrustations of algae are found on rocky shores, notably in Ascension, where coralline algae coat bare rock and boulders. Emerged beaches and shore terraces are found at various levels, indicating episodes of tectonic uplift as well as sea level changes. Many of the islands have beaches of dark volcanic sand, as well as gravel and boulders derived from the

v­ olcanic rocks. Locally there is pale calcareous sand, mainly from shelly organisms, as on Ascension. The Bahamas and Bermuda have coral, limestones and biogenic carbonate sands, and exposures of beach rock. Sable Island, 300 km SE of Nova Scotia, consists of ridges and shoals of quartzose sand derived from glacial drift deposits. Quartzose sands in the eastern Canary Islands may have come from the Sahara. The Atlantic Ocean encompasses climatic zones from Arctic in the north through the mid-latitude westerly wind belt to a calm high pressure zone around 30° north, then the NE trade winds, the equatorial calms, the SE trade winds, another calm zone around 30° south, and the southern mid-latitude westerlies extending into the Antarctic. There is ice on Jan Mayen Island and in Iceland, and on  Bouvet Island in the far south, and frost shattering occurs on rock outcrops on these islands and in the South Shetlands. There are sectors of peat cliff in Iceland and the Falkland Islands. Sea ice forms on the shores of islands in high latitudes. Within the tropics the islands are generally wet and well vegetated, but in the eastern Atlantic the Cape Verde Islands, like the subtropical Canary Islands to the north, are relatively dry. Mangroves are found on the Bahamas and at their northern limit in Bermuda, but there are few niches for mangroves or salt marshes on the rocky volcanic islands. Ocean swell, initiated in stormy areas in the South­ ern  Ocean moves across the Atlantic from the south, arrives as a SW swell on the west coasts of Europe and  Africa, and as a SE swell on the east coasts of the Americas. Storms in the North Atlantic generate waves that arrive as  a NW swell on the west coast of Europe, occasionally  penetrating to NW Africa, and as a NE swell  on the  east  coast of North America, occasionally penetrating  to  Brazil. There are contrasts between the bold coasts of  islands exposed to strong wave action over  long fetches  and those that are more sheltered, ­sometimes intricate and  low-lying. Mean spring tide ranges are ­relatively large  in >Greenland and >Iceland but generally small (Other Atlantic Ocean Islands.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

6.0

Atlantic Ocean Islands – Editorial Introduction

⊡⊡ Fig. 6.0.1 Atlantic Ocean Islands. (Courtesy Geostudies.) Iceland Greenland

Faeroe Is. St. Kilda Rockall

British Is.

EUROPE NORTH AMERICA

RI

DG

E

Sable I.

Azores

Bermuda

IC

Madeira Selvagens

NT

Bahamas

- ATLA

Canary Is.

MID

Cape Verde Is.

AFRICA

Fernando Po Principe Sao Tome Annobon

St. Peter and St. Paul Rocks Atol das Rocas

Ascension

SOUTH

AMERICA

RIDGE

Fernando de Noronha

St. Helena

C

Martin Vaz

TLA

NTI

Trindade

M I D- A

312

Tristan da Cunha

N

Gough I.

4000m. Isobath Falkland Is.

SCOTIA

ATLANTIC OCEAN ISLANDS 0

1000

South Shetland Is.

South Georgia

SEA

Bouvet I.

South Sandwich Is. South Orkney Is.

2000 km

Reference Hansom JD (2005) Atlantic Ocean islands, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 88–95

ANTARCTICA

6.1 Bahamas

Alan K. Craig

1.  Introduction

there are larger fluctuations in sea level related to variations in atmospheric pressure and strong winds. The islands are situated upon the Bahama Banks (>Fig. 6.1.1), large sedimentary shoals typically no deeper than 4 m at low tide, but plunging at their margins to oceanic depths. The four principal Banks (Little Bahama, Great Bahama, Cay Sal, and Acklins) constitute a group that extends to the southeast where, although not politically part of the Bahamas, the Turks and Caicos islands and three more shallow banks (Mouchoir, Silver, and Navidad) to the southeast are geologically similar. They consist of a thick sequence of Cretaceous, Tertiary and Quaternary sediment including calcareous sandstones, limestones, coralline and other biogenic material, accumulated since

The Commonwealth of the Bahamas, east of Florida and north of Cuba, consists of 29 substantial islands and over a thousand cays, islets and rocks with a total area of over 13,000 sq. km and a 3,542 km coastline. They straddle the Tropic of Cancer, and have a subtropical climate with warm summers and mild winters; Nassau has a mean monthly temperature of 21.7°C in January, rising to 27.2°C in July, and an average annual rainfall of 1,179 mm. The prevailing winds are the northeast trade winds, and hurricanes may occur between July and November, passing westward into the Caribbean. Atlantic ocean swell and storm waves arrive from N to NE. The tide range is small, about a metre, but

⊡⊡ Fig. 6.1.1 The Bahamas, Turks and Caicos Islands. 80°

73°

27°

27°

BAHAMA GRAND BAHAMA ISLAND

GREAT ABACO ISLAND

I S L A N D S

ELEUTHERA ISLAND 22°

NEW PROVIDENCE ISLAND TONGUE ANDROS ISLAND

69°

CAICOS ISLAND

EXUMA

OF THE

SOUND

CAT ISLAND

TURKS ISLAND

OCEAN

MOUCHOIR BANK SAN SALVADOR

21° SILVER BANK

CAY SAL LONG ISLAND

GREAT BAHAMA BANK

NAVIDAD BANK 20° MAYAGUANA ISLAND

72°

ACKLINS ISLAND

GREAT INAGUA ISLAND 21° 80°

0

5

73°

10 kms

21°

OIB. L BEFORE JUN. 31

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

69°

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Bahamas

Cretaceous times and resting upon a Jurassic basement which has subsided to a depth of 6,000–7,000 m (Newell and Rigby 1957). They are not coral reefs, but sedimentary structures (notably dune calcarenites that formed when sea level was relatively low) bearing extensive corals in patterns that show variations related to wave energy. They are areas of oolite formation, together with grapestones (aggregates), biogenic sands and silts, and oozes. The sea floor has planed-off coral limestone and dune calcarenite, and ridges and ripples of varying dimensions and mobility, with extensive seagrass growth. The Bahama Banks are trenched by some of the world’s deepest canyons, with steep walls and floors that descend to abyssal plains that may be more than 6,000 m deep (Andrews et al. 1970). These include the wide Tongue of the Ocean, beside Andros Island, and Great Abaco Canyon. There are also circular blue holes, submerged dolines up to 80 m wide, which in some cases emit fresh water from submarine springs. The islands lie near the northeastern margins of the Bahama Banks, the largest being Andros (168 km by 64 km), the others generally long and narrow, aligned at right angles to the dominant northeast waves and often recurved at the ends. They have relatively subdued, riverless landscapes of stabilised calcarenite dunes that rise to a maximum height of 67 m at Mt. Alverina on Cat Island (>Fig. 6.1.1). In general, the northeastern coasts rise relatively steep to a dune calcarenite ridge, followed by a gentler southwestern slope down to lagoons and mangrove swamps. Pleisto­ cene dune calcarenites, generally well indurated (Bahamas Limestone) and having conspicuous karstic features, are common in the interior and underlie dunes up to 20 m high (Hearty and Kindler 1997). The coastlines are dominated by low cliffs and shore platforms cut in Pleistocene aeolian calcarenites that formed as successive, superimposed dune formations during interglacial phases, when the sea was at or close to its present level. These dunes were subsequently stabilised and lithified during glacial phases when the sea was relatively low and the Bahama Banks emerged, and soils (now seen as interbedded palaeosols in  cliff sections) formed, incorporating airborne dust. The dune calcarenites were then trimmed back as cliffs during the Holocene marine transgression and ensuing stillstand. The cliffs and shore outcrops show strong karstic weathering, with jagged pinnacles and rough pitted rock surfaces, as well as notches and visors at the cliff base. Locally there are vertical soil pipes with indurated rims of travertine, which protrude, or have been excavated as banana holes. The islands have windward (NE) and leeward (SW) coastlines, the strongest wave action being from the

ocean to the northeast. Coastlines facing the open Atlantic are subject to large swell and occasional storm waves that sweep in over the narrow coral-covered shelf slope and break on the barrier reef rim. Between the barrier reef rim and the beach there is usually a comparatively broad, shallow lagoon with varying amounts of biogenic sand being transported toward the coastline (Rankey 2002). Microa­tolls occur here, oval or circular structures with diameters of up to 30 m, rimmed by vigorously growing corals and algae. There are also algal mats (stromatolites), thin sheets of blue-green algae, diatoms and other organisms, which trap and bind sediment and induce carbonate precipitation. The beaches are of calcareous sand and gravel, including coralline, algal and oolitic sediment, with occasional outcrops of beach rock. On Cat Island there are sub-parallel Holo­ cene sandy beach ridges up to 5 m high, occurring in truncated sets on successively higher terraces that are occasionally capped with stabilised dunes. Leeward coastlines have wide shelves (e.g., Great Abaco, Andros, Eleuthera, Aklins and Caicos islands) on which highly mobile, fine-grained calcareous sediment occur as shoals. The north coast of Acklins Island has a chain of low narrow barrier islands of caronate sand banked up along the outer edge of shoals. Oolitic sands and grapestone are forming. Large sand shoals are often dissected by tidal channels. Surf reaching leeward beaches is seldom more than a metre high, so that both depositional and erosional features tend to be subdued. Fossil beaches and hypersaline lagoons with abandoned tidal channels are common. Sets of parallel and truncated beach ridges occupy elevations equivalent to the more extensive series found on the windward shore. Emerged Pleistocene marine terraces have been described from the island of Eleuthera (Kindler and Hearty 2000). Some have explained the beach ridges as the outcome of eustatic fluctuations of sea level, each ridge formed during a phase of emergence and trimmed back when the sea returned to its former level (Lind 1968). Others consider that the beach ridges are due to ll major storms alternating with calmer phases (Perkins and Enos 1968; Shinn et al. 1969), particularly on Andros Island. The effects of severe storms on the Great Bahama Bank and shore of the Bahamas in general appear to depend in large measure on the local conditions associated with each event. Some hurricanes have crossed the Bahamas leaving little trace, but most have had major effects on coastlines, eroding beaches and cutting back cliffs. Deep, sub-parallel channels cut through low-lying islands are known as bogues (as on the east end of Grand

Bahamas

Bahama and the centre of Andros), and are believed to be storm-related. Erosion has been a problem on several islands in the Bahamas. In the Turks and Caicos Islands, the former Back Street of Cockburn Town is now Front Street because of coastline recession. Throughout the Bahamas, vegetation has been disturbed by human activities. Nearshore species are halophytic and the Australian pine (Casuarina equisetifolia) is ubiquitous. The interior of Andros Island supports native pine stands, but a complex, dense coppice is typical of other islands. Red mangroves (Rhizophora mangle) are often found along leeward shores.

References Andrews JE, Shepard RP, Hurley RJ (1970) Great Bahama Canyon. Geol Soc Am Bull 81:1061–1108

6.1

Hearty PJ, Kindler P (1997) The stratigraphy and surficial geology of New  Providence and surrounding islands, Bahamas. J Coastal Res 13:798–812 Kindler P, Hearty PJ (2000) Elevated marine terraces from Eleuthera (Bahamas). Glob Planet Change 24(1):41–58 Lind AO (1968) Coastal Landforms of Cat Island, Bahamas: A Study of Holocene Accretionary Topography and Sea-Level Change. University of Wisconsin Newell ND, Rigby JK (1957) Geological Studies on the Great Bahama Bank. Society of Economic Paleontologists and Mineralogists, Special Publication 5:15–72 Perkins RD, Enos P (1968) Hurricane Betsy in the Florida-Bahama Areageologic effects and comparison with Hurricane Donna. J Geol 76:710–717 Rankey EC (2002) Spatial patterns of sediment accumulation on a Holocene carbonate tidal flat, NW Andros Island, Bahamas. J Sediment Res A 72:591–601 Shinn EA, Lloyd RM, Ginsburg RN (1969) Anatomy of a modern carbonate tidal-flat, Andros Island, Bahamas. J Sed Petrol 39:1202–1228

315

6.2 Bermuda

Eric Bird

Introduction Bermuda consists of a group of about 150 small limestone islands in the west Atlantic (32°20' N, 64° 45' W), just over 1,000 km east of Cape Hatteras. They are situated near the south-eastern margin of the submerged Bermuda Platform on part of a narrow submerged reef that encircles a lagoon up to 16 m deep. They have a total area of 54 km2 and are in the form of an italic J, with a 103 km coastline, the longest segment of which faces SSE. The Main Island is 36 km long and generally 2–3 km wide, running WSW-ENE, with a curving southwestern extension round to the Ireland islands, sheltering Great Sound and Little Sound, a recurved inner ridge on which the little city of Hamilton stands, and narrow strips enclosing three large lagoons, Harrington Sound, Castle Harbour and St George’s Har­ bour, at the northeastern end. Fort St. Catherine stands on the northernmost point. There are many sites of scientific interest around the island coasts, but visitors should be warned that some are on private land and inaccessible. The climate is subtropical, under the influence of the warm Gulf Stream, with an average annual rainfall of 1,450 mm. The prevailing winds are westerly, sometimes stormy in winter. In addition to waves generated by westerly winds there is a southerly Atlantic Ocean swell which becomes southeasterly as it breaks on the coast of Main Island. The tide range is small, about a metre. The islands are hilly, rising to summits 60–70 m above sea level, and consist of Pleistocene dune calcarenites (aeolianites) and associated beach calcarenites (Sayles 1931) on the rim of the Bermuda Platform, a submerged seamount truncating a volcano that was last active in the Oligocene. Borings have shown that the base of the calcarenites rests upon an eroded foundation of volcanic lava undulating between 20 and 170 m below sea level. The rock sequence indicates that there has been tectonic subsidence, but the Bermuda Platform is thought to have become stable in Pleistocene times. The dune calcarenites formed from the deposition of successive, superimposed dunes deposited during interglacial phases behind and over Pleistocene beaches that existed when the sea stood at or near its present level (Vacher 1973). These dunes were subsequently stabilised

and lithified during glacial phases, when extensive areas of the Bermuda Platform emerged. Terra rossa soils then formed on the dunes, incorporating airborne dust presumably from North America, and these are now seen as interbedded palaeosols, sometimes associated with indurated calcrete horizons. The rock formations are highly permeable, and there are no rivers. The Pleistocene dune calcarenites were trimmed back as low cliffs and shore platforms during the Holocene marine transgression and the ensuing stillstand. The cliffs and shore outcrops show strong karstic weathering, with jagged pinnacles and rough pitted rock surfaces, as well as notches and visors at the cliff base. Locally there are vertical soil pipes with indurated rims of travertine. The cliffs are fronted by erosional shore platforms that locally carry corals and algae, and associated calcareous sandy beaches. There is evidence of Pleistocene marine terraces up to 22 m above present sea level, but no indication that sea level rose significantly above present datum in the Holocene (Hearty and Kindler 1995). Beaches are calcareous, consisting of calcarenite sand and gravel with some coralline debris. They have a pink colouration, attributed to the presence of red foraminifera (Homotrema rubrum). Sand and gravel from the bordering reefs is still being delivered to beaches, but on the northern (lagoon) shore the supply has diminished and the coast is cliffed with only a few pocket beaches. The cliffed coastline has numerous small bays and coves formed by marine submergence of low-lying parts of the dune calcarenite landscape. Sinky Bay is an almost circular cove. There are caves, natural arches and stacks. Mangrove Lake is the largest of five shallow brackish (anchialine) ponds with limited tidal flushing through fissures and small caves. Coral reefs are extensive off the island coasts, those on the northwest side being the northernmost reefs in the Atlantic Ocean. They have been damaged by the dredging of navigation channels and the grounding of ships. Dundonald Channel is a straight navigation channel cut across a broad shoal to access Great Sound and Hamilton Harbour. Microatolls are found, oval or circular structures with diameters of up to 30 m, rimmed by vigorously growing corals and algae, and there are also serpulid ridges on

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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6. 2

Bermuda

calcarenite shore platforms and cup reefs built by vermetid gastropods and algae. Sinkholes formed by solution during low sea level phases persist in coral reefs and platforms of calcarenite. Mangroves grow at their northernmost limit (32° N) in  Bermuda, and include the red mangrove (Rhizophora mangle), black mangrove (Avicennia germinans) and button mangrove (Conocarpus erectus). They have responded to a sea level rise of 28 cm over the past century by aggrading a peaty substrate, but there has been recession of their seaward margins (Ellison 1993). 

References Ellison JC (1993) Mangrove retreat with rising sea level, Bermuda. Estuar Coast Shelf Sci 37:75–87 Hearty PJ, Kindler P (1995) Sea level highstand chronology from stable carbonate platforms (Bermuda and the Bahamas). J Coastal Res 11:675–689 Sayles RW (1931) Bermuda during the Ice Age. Proc Am Acad Arts Sci 66:381–468 Vacher L (1973) Coastal dunes of younger Bermuda. In: Coates DR (ed) Coastal geomorphology. State University of New York, Binghampton, USA, pp 355–391

6.3 Greenland

Introduction The coast of Greenland is related to the variety of rock types, and the activity of glacier ice during the Pleistocene and Holocene. Much of the western and southeastern coast consists of bedrock, often fringed by sea ice. The coast facing the Polar Sea is primarily built up of PreCambrian sedimentary rock, whereas the northeast coast is dominated by Palaeozoic and Mesozoic formations. Younger geological strata, such as erodible Cretaceous sandstone and plateau basalts of Tertiary age are found both on the east and the west coast, but only between latitudes 69° and 72° N (Escher and Watt 1976). Most of the rocks outcropping along the Greenland coasts are very resistant to marine erosion, and the coastal landforms bear the strong imprint of glaciation. The ice cap has been much more extensive, covering nearly all of Greenland until about 12,000 years ago. The effects of glaciation are most obvious in the numerous large fiords, where the ice cap, through selective erosion along preexisting (Tertiary) valleys, created deep troughs that were invaded by the rising sea as the ice cap partially melted. More broadly, the spreading ice smoothed rock ridges and dissected coastal outcrops into scattered hills, which with postglacial marine submergence have become skerries on archipelagic coasts. Greenland covers an area of about 2.2 million km2, of which only 342,700 km2 are ice-free. The precise length of the coastline is unknown, but is estimated to be at least 40,000 km. The range of variation in coastal geomorphology is nevertheless relatively limited. About 90% of the coast is rocky, with more than 100 long, deep and much branched fiords. In Scoresby Sound on the east coast there are 11 fiords with lengths as great as 325 km and depths of up to 1,450 m. Present-day marine erosion and accumulation are not conspicuous, but there are gravelly pocket beaches (pebbles to boulders), and on basalt and limestone outcrops marine processes have shaped coastal landforms. On the west coast island of Disko, for example, there are high

cliffs along the south and west (> Fig. 6.3.1) and long shingle barriers with lagoons and marshes on the east (Nielsen 1969). On parts of the Greenland coast the ice cap still reaches the sea. In Melville Bay, northwest Greenland, an ice front forms cliffs up to 40 m high almost 500 km and farther north, at Kane Basin, ice cliffs extend for about 100 km. Many valley glaciers discharge into the sea, some of them producing icebergs. When annual mean wave energy levels are calculated for the outer coast on the basis of wind statistics and fetch, far higher values are found than those that actually occur. This is partly due to the freezing of the sea, and partly to the arrival of ice floes from the Polar Sea. These pass down the east coast, round the southernmost point, then northward, carried by the Irminger Current along the west coast to about 65° N latitude. There are also many icebergs from the calving glaciers. All these types of ice hamper the transmission of waves to the coast. The sheltering effect of the ice favours various accretion forms shaped under low wave energy conditions. Where there is occasionally no ice, the full energy of the ocean waves reaches the shore. Nanok (ice bear) hunters on the northeast coast have reported that a single gale completely changed the configuration of the coast, which had been stable as long as could be remembered. In this case the Storis, the constantly south-flowing ice-laden current from the Polar Sea, had been blocked for some time by heavy ice farther north, and thus could not exert its usual subduing effect on the waves. Although this situation is rare, such events must also be considered when interpreting coastal geomorphology. Mean spring tide range varies from about 3 m in the southwest at Nanortalik (60° N), 4.5 m at Godthab (64° N), 2.5 m at Disko Island (70° N), to 3.5 m at North Star Bay (77° N). The coasts facing the Polar Sea have almost no tide, only 0.3 m being recorded at Kap Morris Jesup (83° 40' N). Along the east coast the mean spring tide is about 3.0 m in the south, decreasing northward, to less than 1.0 m at Nordostrundingen, the northeastern corner of Greenland (81° N).

Edited version of a chapter published by Niels Nielsen in The World’s Coastline (1985: 261–265). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Greenland

⊡⊡ Fig. 6.3.1 Cliffs cut in a raised marine terrace. Cirques with rock glaciers are common along the west, north, and northeast coasts of Disko Island, and in many places the rock glacier front forms an advancing cliff. The upper edge of the plateau basalt lies about 900 m above sea level.

When the sea freezes in winter the tide is of great significance because of ice foot formation. On sedimentary coasts this leads to the formation of thermokarst on the backshore, and accumulation or erosion when parts of the ice foot, containing gravelly material, are driven up on the shore or rafted away from it. These processes result in the displacement structures characteristic of arctic beaches. Erratic rocks and alien sediment incorporated in the ice foot are stranded on the shore when it melts (Nielsen 1978). The ice foot may also influence rocky coasts by breaking off rock fragments in the springtime, when the warmer sea melts the ice foot from below, so that it collapses into the sea (Nielsen 1979). Occasional large waves are produced suddenly by calving ice cliffs or capsizing icebergs. Though rare, this type of wave is both high and long, and can result in the deposition of wave-transported sediment up to 15 m above sea level in a fiord where the fetch is very limited (Nielsen 1969). Throughout most of the Quaternary period, the ice cap covered Greenland, though varying in extent. In Late Pleistocene times, about 12,000 years ago, when the climate began to become warmer, the ice margin receded, sea levels were rising, and the reduction in ice mass resulted in rapid isostatic uplift in coastal regions. This relative upheaval of the land has led to the emergence of Late Pleistocene and Holocene coastlines, seen as raised marine terraces along many parts of the Greenland coast (> Fig. 6.3.2). In East and West Greenland the isostatic upheaval of land started about 9,000 years ago and apparently ended 4,000–5,000 years later (Hjort 1979). A similar development seems to have taken place in North Greenland,

although somewhat delayed by the later melting of the ice cap in this part of the country. Thus the formation of coastlines that are now about 10 m above sea level, has been dated to 3,000–4,000 years ago (Weidick 1972; Stablein 1975). Emerged coastal features include marine terraces and beach ridges that reach elevations of about 60 m in South Greenland, compared with 100–200 m in the central part of West Greenland, and up to 216 m above present sea level in East Greenland, indicating the pattern of the highest Holocene marine limit. The upheaval of land has varied from one region to the other because of local variations during the melting phase. Minor fluctuations in sea level relative to the land are primarily related to eustatic changes in sea level (Humlum 1980). After the Little Ice Age, the climatic deterioration that set in during the seventeenth century, a general sea level rise of about 1 m occurred, resulting in the submergence of ruined buildings. During the past 50 years there appears to have been a slight lowering of sea level relative to the land, but this may now have reversed. Ammassalik Island off SE Greenland is dominated by high ground, presumably upwarped along the contintental margin, with a narrow coastal plain. During the Quaternary the island was covered by the Greenland Ice Sheet, with the exception of a few outcrops on the escarpment edge. Periglacial processes have also been active, intensifying during the late Holocene (Humlum and Christtansen 2008). In Ikkafjord in southwestern Greenland are the so-called Viking Warriors, towers of tufa that grow up from the shallow floor of the fiord to a height of 20 m,

Greenland

6. 3

⊡⊡ Fig. 6.3.2 Raised marine terraces in Mellemfiord, west Disko. The border between the scree from the basalt scarp and the upper terrace edge lies 53 m above sea level, and was formed about 9,000 years ago.

occasionally protruding from the water on calm days. They can grow upward 0.5 m in a year, and have formed as the result of exudations of alkaline groundwater from sea floor fissures. Minerals dissolved in the groundwater crystallise into a mineral called ikaite, which is added to the top of the growing towers. The feature is geologically unique.

References Escher A, Watt WS (eds) (1976) Geology of Greenland. Geological Survey of Greenland Copenhagen, Denmark Hjort C (1979) Glaciation in northern East Greenland during the late Weichselian and Early Flandrian. Boreas 8:281–296

Humlum O (1980) Arctic geomorphology, Enoks havn, Disko (in Danish). Geogr Noter 5:1–142 Humlum O, Christtansen HH (2008) Geomorphology of Ammassalik Island, SE Greenland. Geografisk Tidsskrift 1089:50–20 Nielsen N (1969) Morphological studies on the eastern coast of Disko, West Greenland. Geografiska Tldsskrift 68:1–35 Nielsen N (1978) Coastal morphology in the Arctic region (in Danish). GO. Geograforlaget Brenderup:242–247 Nielsen N (1979) Ice-foot processes. Observations of erosion on a rocky coast, Disko, West Greenland. Zeitschrift für Geomorphologie, N.F. 23:321–331 Stablein G (1975) Ice-margin deposits and coastal evolution in West Greenland (in German). Polarforschung 45(2) Weidick A (1972) Holocene shorelines and glacial stages in Greenland – an attempt at correlation. Geological Survey of Greenland 41

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6.4 Iceland

Eggert Larusson

1.  Introduction Situated in the North Atlantic Ocean between 63° 30' N and the Arctic Circle (66° 30' N), Iceland is a volcanic island with an area of 103,100 sq km. It is the outcome of Tertiary and Quaternary extrusions and intrusions of volcanic rocks (lavas, tuffs, pyroclastic sediment) from the mid-Atlantic Ridge, a zone of crustal plate divergence. It has been glaciated, and there are five residual ice areas. Glacial drift deposits are found widely, particularly bordering the residual glaciers, where terminal moraines in­­ dicate recession of ice margins in recent decades. Much of the coastline is irregular, with fiords (known in Iceland as fjördur) between mountain ranges, but on the southern coast between the Olfusa estuary and Djupivogur, large volumes of sand and gravel deposited by powerful streams fed by melting glaciers have formed a wide lowland (sandur), with a smooth coastline of wide bays and broad lobate promontories (>Fig. 6.4.1). The lowland is backed by cliffs and bluffs marking a 9,000 year old cliffy coastline that existed before this extensive deposition occurred. The landscape of Iceland is dominated by basalt plateaux and glaciated ridges, with 10% still covered by ice. Emerged beaches are found on many parts of the Icelandic coast (Jonsson 1957; Norddahl and Petursson 2005). The height of the Postglacial marine transgression limit differs from place to place: 130–140 m in the west, but only 35–50 m in the north, the east, and the south. The highest Holocene coastlines have been dated at 12,300–9,600 bp. The sea reached about its present level by 9,000 bp and the isostatic uplift was complete 3,000 years later. In the interim, the Holocene marine transgression was more or less matched by the isostatic recovery. Before 3,000 bp, sea level seems to have stood about 2–4 m below the present level, and a small further transgression may then have taken place (Thorarinsson 1956). Roches moutonnées, striations and some till cover indicate that the Icelandic strandflats existed before the last glaciation. The climate is cool temperate oceanic. Reykjavik, the capital city on the southwest coast, has mean monthly temperatures of 0°C in January and 11°C in July, with an average annual rainfall of 800 mm. Affected by the ­numerous

arctic (polar front) cyclones, Iceland is windy. Wind speed and direction are very variable, but the most frequent winds blow between southeast and northeast, and the main track of the depressions passing close to the southern coast. Gale force winds are common. There are extensive areas of bare, unvegetated landscape, rocky or gravelly, as the result of strong winds, low temperatures and desiccation. Peat deposits are of limited extent, mainly around lakes and lagoons. Some shore ice forms in cold winters, especially in the northern fiords. The coasts of NW Iceland are less stormy than those of the south and east coasts. Prevailing winds and waves in coastal waters are from SE and NE rather than SW, while the southern coasts receive an Atlantic ocean swell from the south. Mean spring tide ranges are generally between 1.5 m and 3 m, and a little more along some parts of the western coasts, where Reykjavik has 3.8 m and Grindavik in the SW has 3.2 m. Bakkagerdi on the NE coast has 1.5 m and Bolungarvik in the NW has 2.4 m.

2.  The Icelandic Coastline The NW peninsula consists of much-dissected Tertiary and volcanic formations (the oldest in Iceland), and presents an indented rocky coastline of deep, steep-sided fjördur. There are very steep coastal slopes and megacliffs (300–600 m) on which landslides sometimes occur. Deep inlets branch from the fiord of Jokulfjördir, Leirufjördur on the southern side receiving glacifluvial sand and gravel from a stream draining out of the Drangajökull glacier. There are several large snow-filled corries and waterfalls below hanging valleys. On the southern side of Isafjadardjup are several parallel deep fiords (fjördur) running from south to north. The Isafjördur is the easternmost, with a grey sandy beach in the bay at the mouth of the Laugadalsa, but much of the fiord coast is rocky with angular gravels in coves. On the western side of Isafjördur is a cuspate spit, flanked by shingle beaches that fade out along the shore (>Fig.  6.4.2). Beaches and spits in these north-western fjördur are related to streams delivering shingle, or to eroding sectors

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

6.4    

Iceland

⊡⊡ Fig. 6.4.1 Some coastal features of Iceland. 1 – glaciers, 2 – subglacial volcanoes and/or hot zones, 3 – outwash plains (sandurs) near the coastline, 4 – track of main glacier bursts (jökulhlaups), 5 – beaches and barrier spits, 6 – predominant longshore drift, 7 – strandflats. (Courtesy Geostudies.) 1 4°

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⊡⊡ Fig. 6.4.2 Cuspate spit on the western shore of the Isafjördur, NW Iceland. (Courtesy Geostudies.)

Iceland  

of glacial drift. There are several cuspate spits, shaped by waves that approach diagonally across these long, narrow inlets. There are sandy beaches and flanking river mouths at the head of each fjördur, but in several sectors the coastline is artificial because of road construction and protective boulder walls. Cliffs are found towards the more exposed northern ends of the fjördur, and are high and bold towards Gardstadir. Skötufjördur, Hestfjördur, Seydisfjördur and Alfdafjör­ dur are parallel fjördur on the north coast. Sudavik stands on a morainic slope behind a beach and a lobate foreland. In January 1995, 22 houses were swept away by an avalanche, and a new town was then built to the south. The next deep inlet is Skufulsfjördur, with high volcanic cliffs and slopes, and an incipient corrie on its eastern slopes. The town of Isafjördur stands on its western side, a port built on a former recurved spit, protected from northerly storms by a large boulder wall. West of Isafjördur, the coast is steep on glacial drift, with screes and slumping slopes. The beach is of dark volcanic sand and gravel with splays of pale brown calcareous sand, backed by low grassy dune ridges, then a scatter of small dune hummocks (>Fig. 6.4.3). Bolungarvik harbour is protected by a large boulder breakwater. Round a high promontory is Sugandafjördur (>Fig. 6.4.4), with the Stadur valley on its southern side. Steep slopes rise to a high plateau, and at the valley mouth is the cove at Skeravik (>Fig. 6.4.5). To the east is Sudureyri, a fishing port. The upper part of Sugandafjördur is shallow, with gravelly shoals exposed at low tide. ⊡⊡ Fig. 6.4.3 The shore at Bolungarvik, NW Iceland. The dark beach is of basaltic sand and gravel, the paler material calcareous sand from the sea floor. (Courtesy Geostudies.)

   6.4

On the north coast of Onundarnfjördur, the town of Flateyri has been built on a spit that curves upfiord. At the head of this fjördur the beach has a covering of brown comminuted shell sand. At Holt is an inner spit, concave seaward, which has a brown calcareous sandy beach bac­ ked by grassy dunes and partly eroded hummocks. Opening to the west coast is Dyrafjördur, with Pingeyri on a sandy cuspate spit, with a brown sandy beach on the southern side. Upstream the beaches and spits are of grey volcanic sand and gravel. Arnarfjördur is steep-sided, with some minor spits, and cliffs over 500 m high at Selardalshidar face WSW into the open ocean. It branches headward into several smaller fjördur, one of which is Dynjandisvogur, where steep coasts in glacial drift slope down to beaches of grey sand and gravel with several cuspate spits. A spit has grown half way across Talknafjördur. In Patreksfjördur, the town of this name stands on two sand spits, Vatneyri and Geirseyri on the northern shore, while Sandoddi is a recurved spit on the southern shore. At the western end of the Bardastrandar Peninsula Blakknes is a 281 m headland and Breidavik has a curved beach. High cliffs run along Bjarnarnupur, and on the southern side of the peninsula the cliffs of Latrabjarg rise to 440 m. Raudisandur is a long barrier enclosing a marshy area, and Baejarvadall is a coastal lagoon. To the east, high cliffy peninsulas separate a succession of fiörden on the northern coast of the broad Breidafjördur. Breidafjördur is a submerged strandflat, with numerous islands, skerries, and shoals. One of the islands is

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⊡⊡ Fig. 6.4.4 Sugandafjördur, NW Iceland. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.5 Stadur, west of Skeravik, looking towards Saudanes. The steep coastal slopes descend from the coastal plateau to an emerged terrace. (Courtesy Geostudies.)

Flatey, with a low cliffy shore where columnar basalts are disintegrating into aprons of basalt boulders. On the east coast of Breidafjördur, a strait narrows to Gilsfjördur, and to the south is the mountainous Klofningsvegur Peninsula, then the Hvammsfjördur strait.

On the Snaefellsnes Peninsula to the south the width of the emerged strandflat is generally 10–20 km, backed by steep bluffs. South of the Hvammsfjördur is the Skog­ arströnd coast, which has a terrace incised by rocky gorges, such as those cut by the Dunkur and Skrauma rivers.

Iceland  

Gravelly watercourses cross the coastal plain, and there are seaweed-strewn shores of rocky lava, and few beaches. An archipelago includes the large embayed island of Brokey. Eyrarfjall is a high volcanic ridge beyond which two valleys converge to the Osa. Alfafjördur is a major inlet, backed by a salt marsh and an emerged fjördur floor ringed by rocky cliffs. Stykkisholmur is a ferry port on a peninsula on the south coast of Breidafjördur. To the west is the coastal district of Helgafellssveit, with a terrace that descends northward to low cliffs and rocky shores with gravel at river mouths. Incised streams tumble over waterfalls to the sea on this emerged coast. There is a wide lowland with an extensive spiky lava field, and inlets with strong tides. Pass Eidi leads westward through to the bay of Grundarfjördur. West of Grundarfjördur is the mountainous peninsula of Kirkjufell (Church Mountain), 463 m high. At Buland­ shöfdi, there are said to be marine shells 40–55 m above sea level. Gamlavik is a broad bay backed by two barrier lagoons separated by a delta. At Ondverdarnes, the coast turns southward along the dark Svörtuloft cliffs, cut in Holocene lava flows. Similar cliffs fringe the western part of the Snaefellsnes Peninsula, which is dominated by blocky lava deposited within the past 11,000 years. On this coast, beaches are found only where there is a nearby source of sand or gravel, either a cliff or a river mouth.

⊡⊡ Fig. 6.4.6 A boulder beach derived from lava outcrops in the cliffs at Hellnar. (Courtesy Geostudies.)

   6.4

At Hellnar, the cliffed coast turns north-eastward. There is a sharp contrast between smooth pastoral land on tuff at Hellnar and the irregular bouldery lava flow to the north. Cliffs cut in tuff come to an end at the harbour, ­giving place to irregular lava cliffs dissected by inlets and a natural arch. The harbour contains a sandy beach, derived from tuff in the adjacent cliffs and on the sea floor. A boulder beach of dark and pale volcanic rock occupies a small bay to the north of the breakwater (>Fig. 6.4.6). East of Arnarstapi is the Midhusavatn lagoon, separated from the sea by the wide sandy Hraunlandarif barrier. Further east a hummocky lava flow descends from the escarpment at Axlarholar, out across the gently sloping coastal lowland to form the lobate foreland of Budarhaun. This is a fine example of a lobate foreland produced by volcanic deposition, and includes a cinder cone. Budir stands at the eastern end of the rugged Budahraun lava field. There is a tidal lagoon, Budaos, and a curved beach of light brown calcareous sand backed by grassy dunes extends eastward past Bardastadir, where an inlet is bordered by low cliffs exposing stratified lagoon deposits that had been overrun by dunes on a receding coastline (>Fig. 6.4.7). Stakkhamar is a gravelly plain, and a chain of sandy barrier islands extends round to Hitarnes. The coastline then turns southward. There are extensive beaches and dune-capped barriers fringing lagoons, some of which are behind wide grassy strips that may include glacial drift. South of the Hafjardara River, the

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⊡⊡ Fig. 6.4.7 Stratified lagoon deposits exposed on the shore at Bardastadir. (Courtesy Geostudies.)

coast is backed by the wide Myrar coastal plain that is like a strandflat, but includes areas of glacial and fluvial deposition. It has recent knobbly or bouldery lava flows, and the Eldborg volcano rises prominently. The coastal plain is interrupted southward by the Borgarfjördur inlet at Borgarnes. On its northern shore, the Langa River cascades over lava ledges, as it drops to the sea. To the south, skerries fringe the low-lying coast past Höfn, and a bold volcanic escarpment overlooks the coastal plain. Hvalfjördur is fringed by strandflats, with steep headwalls. South of Hvalfjördur the Kjalarnes Peninsula is backed by a volcanic escarpment with screes and gullies that feed into glacifluvial channels incised across a coastal lowland that ends in a low cliffy coastline. The hilly embayed Reykjavik Peninsula is urbanised and bordered by sea walls and boulder ramparts, and a large harbour. There are outlying grassy low islands. South of Reykjavik is the Keflavin Peninsula, a grey mossy and grassy Holocene plain with volcanic cones and dykes, and low rocky shores. This is a continuation of the mid-Atlantic Ridge and is consequently a young volcanic area, with lavas deposited in the past 11,000 years. The southern coast is exposed to strong Atlantic swell and storm waves. Low cliffs cut in Holocene lava flows predominate, and high level storm beaches of loose boulders and pebbles that have been piled up on top of some of the cliffs by the very large waves that break upon this coast in winter storms. Beaches of well-rounded cobbles and

boulders dominate the shore. Sandy beaches are uncommon, but small isolated dune fields can be found. To the east, a coastal plain backed by a line of cliffs and bluffs marking an early Holocene coastline. Several rivers, some fed by hinterland glaciers, flow southward across a sand and gravel outwash coastal plain (Hine and Boothroyd 1978). From the Pjorsa River a straight sandur coast extends eastward to the large lobate foreland in the lee of the Vestmannaeyjar islands. Southwest of this archipelago a new volcanic island, Surtsey, began to grow in 1963. After the volcanic activity came to an end the island was re-shaped by storm waves which cut back cliffs and deposited beaches of derived sand and gravel. Recession of the cliffs was most rapid on the exposed southwest coast, whereas a cuspate foreland was built on the more sheltered northern coast by convergent longshore drifting (Norrman 1980). The coastal plain, backed by lava cliffs, continues eastward, and the old coastline becomes a grassy slope up to the edge of the Eyjafjallajökull glacier, which lies over a volcano that erupted in 1612 and 1821–1822, causing massive ‘glacier burst’ flooding and the deposition of the wide sandur plain. The 60 m high Skogafoss waterfall pours over the cliff into the River Skoga. The coast is fringed by gravel beaches with some dunes (Mountney and Russell 2006). At Dyrholaey, a spur of high ground runs out to form a 110 m high cliffed headland, with a natural arch. Behind Dyrholaey, a lagoon is impounded by a barrier spit, and a

Iceland  

high rocky ridge of Upper Pleistocene pyroclastic and tuffaceous sediment (Reynisfjall) runs south to a prominent 200 m high headland. On the next headland there is a large cave in columnar lava flows (>Fig. 6.4.8) and The Needles (Reynisdrangar) are rocky stacks 66 m high close inshore (>Fig. 6.4.9). A high scarp then runs inland besides Vik.

⊡⊡ Fig. 6.4.8 Columnar lava on the coast at Dyrholaey. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.9 The Needles (Reynisdrangar), stacks off the cliffed coast near Vik. (Courtesy Geostudies.)

   6.4

Vik is not a coastal town, and there are no harbours along this coast. At Kerlingardalur there are hummocky dunes on the coastal plain, behind a wide black sandy beach (>Fig. 6.4.10). East of Vik is the wide sand and gravel plain of Myrdalssandur, with lava blisters, extending to the braided

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⊡⊡ Fig. 6.4.10 Coastal dunes east of Vik. (Courtesy Geostudies.)

estuary of the Kudafljot River, which is fed by meltwater from the Myrdalsjökull glacier. Eruptions of Katla have occurred every 40–80 years, melting the Myrdalsjökul glacier and sending catastrophic floods southward over the Myrdalsandur. The last of this glacier burst floods (jökulhlaups), occurred in 1918, so another glacier burst is overdue. The bluffs behind the coastal plain decline very gradually beneath lagoons behind barriers, as at the 76 m high island of Ingolfshöfdi (King 1956), passing through broad wet sandy flats with a wide zone of low mounds of windblown sand. The lagoon is more than 6 km wide but only up to half a metre deep. Beaches of dark volcanic sand alternate with zones or ridges of cobbles. Predominant longshore drifting diverges westward and north-eastward from Ingolfshöfdi. On the east coast near the Gljufura canyon there are terraces and barrier-enclosed lagoons, and to landward, is a scarp capped by the glacier lobe of Breidamerkurjökull, part of the huge Vatnajökull ice cap. To the north, the coastal lowland has ridges of glacial drift backed by a glacier lake, Fjallslon, with a clear ice front of the Fjallsjökull lobe, several icebergs and a steep sided cliffed moraine of gravel and boulders. At the Jökulsa River, another morainic ridge encloses Jökulsarlön, a lake crowded with icebergs. The Breidamerkurjökull glacier is melting into a coastal lagoon (>Fig.  6.4.11), which has an outlet to the sea through the sand and gravel barrier. Comparison of maps made in 1945, 1965 and 1998 show that the glacier front has receded by up to 3 km (Evans and Twigg 2000).

The coast of Skeidararsandur has few beach ridges, no tidal inlets, no estuaries, no true lagoons, and no marshes, because of the high rates of fluvial and shore sedimentation. The so called Hornafjördur, Skardsfjördur, Papafjördur, Lonsfjördur, Alftafjördur and Amarsfjördur are shallow lagoons, not true fiords, surrounded by low lying land (strandflats and sandurs). Large volumes of material are deposited in the lagoons, and glacifluvial deltas are presently growing. Tidal inlets are open, and narrow marshes and tidal flats fringe parts of the lagoons. Between Skardsfjördur and Papafjördur, a high rocky promontory of Tertiary lava interrupts the barrier beaches. Another such mountainous sector of steep rocky coast with a narrow strandflat extends between Lonsfjördur and Alftafjördur. North of Djupivogur is a typical fiord coast, with some irregular strandflats, as at erufjördur. Stodvarfjördur, Fastrudsfjördur and Reydarfjördur are fiords at the mouths of glaciated valleys (>Fig. 6.4.12), with intervening ridges ending in cliffy promontories. More fiords are seen north to Glettinganes, where the steep coast turns NW. Heradsfloi, a bay on the NE coast, is a typical sandur fed by rivers draining north from the Vatnajökull glacier. Beyond Vopnafjördur, the northeast coast has lower promontories and wider bays and the narrow Langanes Peninsula declines to Fontur, the northeast point of Iceland. The northern coast of Melrakkasletta is very irregular, with a succession of promontories and bays with lagoons and lakes separated from the open sea by wide barrier embankments of boulders.

Iceland  

   6.4

⊡⊡ Fig. 6.4.11 Breidamerkurjökull glacier calving into a coastal lagoon near Jökulsa. Breidamerkurjökull. (Courtesy Geostudies.)

⊡⊡ Fig. 6.4.12 Glaciated valley ­descending to Mjoifjördur on the east coast of Iceland. (Courtesy Tony Stutterd.)

South of the wide bay of Axarfjördur, a broad sandur is occasionally reached by glacier-burst floods, and some old barrier beaches indicate that there has been coastal progradation. Eyjafjördur is a deep fiord and on the east coast of Skagafjördur is the double tombolo of Thordarhöfdi, linking an island over 200 m high (Bodéré 1973). At the southern end of Skagafjördur barrier beaches and spits block river mouths on either side of the Hegraxes ridge, which protrudes northward.

Beyond the next promontory is the wide bay of Hunafjördur, the southern part of which is split by a protruding steep-sided ridge into Hunafjördur to the east and Hrutafjördur to the west. Hunafjördur has a narrow 7 km long and up to 5 km wide sand barrier (Pingeyrasandur) that divides the Hop lagoon (Hunafloi). At the western end of this beach is the strange stack of Hvitserkur, the remains of a volcanic neck pierced by natural arches and daubed with white gauno.

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Hrutafjördur, by contrast, is a long fiord that narrows southward to a steep-sided valley at Melar. On its west coast are several large tributary fiords, notably Stein­ grimsfjördur and Reykjarfjördur, then the mountainous promontory at Krossnesfjall. To the north–west the fiords are smaller between sharp, narrow peninsulas along the coast to Hornbjarg, where the north-facing cliffs rise to more than 550 m.

References Bodéré JC (1973) Le tombolo double de Thordarhöfdi. Norois 78:213–235 Evans DJA, Twigg DR (2000) Breidamerkurjökull 1998 map (1: 30,000). University of Glasgow and Loughborough University, UK

Hine AC, Boothroyd JC (1978) Morphology, processes, and recent sedimentary history of a glacial outwash plain shoreline, southern Iceland. J Sediment Petrol 48:901–920 Jonsson J (1957) Notes on changes of sea level in Iceland. Geogr Ann 39:143–212 King CAM (1956) The coast of south east Iceland near Ingolfshöfdi. Geogr J 122:241–246 Mountney NP, Russell AJ (2006) Coastal aeolian dune development, Sólheimasandur, southern Iceland. Sediment Geol 192:167–181 Norddahl H, Petursson HG (2005) Relative sea level changes in Iceland: new aspects of the Weichselian glaciation of Iceland. In: Caseldine C et al (eds) Iceland: modern processes and past environments. Elsevier, Amsterdam, the Netherlands, pp 25–78 Norrman JO (1980) Coastal erosion and slope development in Surtsey Island, Iceland. Z Geomorphol 34:20–38 Thorarinsson S (1956) The thousand years struggle against ice and fire. Bokautgafa Menningarsjods, Reykjavik, Iceland

6.5 Other Atlantic Ocean Islands

Jim Hansom

1.  Introduction

2.1.  The Azores

A succinct review of the Geomorphology of Atlantic Ocean Islands, including > Bahamas, > Bermuda, > Greenland, > Iceland can be found in Hansom (2005). The Atlantic Ocean islands fall into three groups (>Fig.  6.5.1): those that lie on the Mid-Atlantic Ridge, from Iceland south to Bouvet Island; those that lie on the American shelves to the west, such as Bermuda and the  Falkland Islands; and those that lie on rises and the Eurafrican Shelf to the east, from the Faeroes south to the islands in the Gulf of Guinea. In stark contrast with the low coral islands and cays of both Indian and Pacific oceans, the Atlantic is mainly characterised by high islands, many of them of volcanic origin. Although extensive in Caribbean waters, in general, coral reefs are poorly developed in the Atlantic Ocean and even the Atol das Rocas, off Brazil, is not a true coral atoll. The Atlantic islands show variations in size and shape, and features that reflect contrasts in their lithology, in their vulcanicity (whether active, dormant or extinct), and in climate and vegetation (the lush tropical Azores to the glaciers of Bouvet Island). Details of the geology and coastal geomorphology of the Atlantic islands are ­discussed in Mitchell-Thome (1970, 1976) and Hansom (2005).

The Azores (Portuguese) (38° 30' N, 28° 00' W – 1,500 km west of Lisbon) consist of nine widely separated mountainous, but fertile islands (total area, 2,344 sq km). The temperate oceanic climate is mild (January 14°C, July 22–33°C), but with abundant rainfall (about 950 mm/year, winter maximum), vegetation cover is extensive. The steep slopes of the islands are dissected by deep and narrow gulleys, but relics of emerged marine beaches and platforms are found at various levels of up to 60 m (Guilcher and Battistini 1982). The coasts are generally steep, with cliffs up to 600 m high, cut into lavas, tuffs, and scoria, sometimes delivering gravel to the shore via scree slopes, but there are also many small beaches of dark volcanic sand. Some white calcareous beaches are also found. The sea is frequently stormy, particularly on north-facing coasts, and the harbour at Ponta Delgada on Såo Miguel is artificially protected by tetrapods. With deep water close inshore (the 100 m isobath is mainly within 3 km of the coastline), the power of the waves reaching the Azores is such that Pico was chosen as the site of a pilot wave-power plant. The coast of the largest island, Sao Miguel, is characteristic of the Azores coastline, in general (>Fig.  6.5.2). The coast is backed by higher volcanic peaks, where steep cliffs fronted by caves, patchy low shore platforms, offshore stacks and islets are found, such as at Mosteiros. Between the cliffed indentations, short sandy beaches occur. Longer sandy beaches occur where the hinterland is lower, such as at Praia dos Moinhos and at Populo, where small pine-clad sand dunes occur. Pico Island is 42 km long and dominated by the cone-shaped stratovolcano of Ponta do Pico whose smoking crater lies at 2,351 m. Various historical lava flows have moved the coastline seaward but have now been eroded back to form low cliffs. There are 600 m high cliffs on the west coast of Flores, which has seven volcanic craters. in sheltered places from the dominant westerly waves smaller volcanic forms survive on the coast, such as at Barca on the island of Graciosa, where a volcanic cone has been eroded into bays fringed by crumbling cliffs of ash that cascade onto a boulder beaches below. The 1957/1958 eruption at Capelhino on Faial showered the adjacent coast

2.  Islands of the Mid-Atlantic Ridge All of the islands of the mid-Atlantic Ridge are of volcanic origin, having been produced by the upwelling of basaltic magma into the crustal gap at the mid-ocean spreading centre in the mid-Atlantic. Many remain active volcanoes, rising from deep water, their coastlines characterised by steep cliffs cut into lava, ash or hyaloclastite. The volcanoes are all youthful and the island coastlines are in a state of continuing adjustment in response to ongoing volcanicity, marine and slope activity. The islands extend south from Iceland and include the Azores, St. Paul and St. Peter Rocks, Ascension, St. Helena, Tristan da Cunha, Gough Island and subantarctic Bouvet Island.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_6.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 6.5.1 Location map.

in ash and contributed to accretion of the beach. At Porto Pim, also on Faial, a substantial sandy spit has developed.

2.2. St. Paul and St. Peter (São Paulo e São Pedro) Rocks More than 800 km from South America, these 12 Brazilian islets (0° 56' N, 29° 22' W) are composed of mylonitic

­ eridotite and are thought to be unaltered mantle matep rial. Dated isotopically at 3,500–4,500 million years, they are among the oldest outcrops on the earth’s surface (Smith et al. 1974), the submarine mountain of which these rocks are the pinnacles extending 4,000 m into the ocean depths. The only source of freshwater on the St. Peter and St. Paul’s Rocks is from rain, and there is almost no vegetation. The lack of vascular vegetation is most likely due to constant sea-spray bombarding the rocks, and the terrestrial

Other Atlantic Ocean Islands

6.5

⊡⊡ Fig. 6.5.2 Stormy weather in the Azores on the north coast of Såo Miguel Island, near Capela. (Courtesy A. Stutterd.)

v­ egetation is dominated by fungus and algae. The rocks bear extensive deposits of white guano: calcium phosphate derived from seabird droppings. The highest point is only 23 m above sea level, and there is an emerged marine platform at 5 m. The coasts are cliffed and rocky, with wavescoured potholes.

Lithothamnion) growing in the narrow shallow water fringing the island. Erosion of adjacent rock shores and terrestrial sediment in-washed by infrequent rains probably account for the remainder of the beach sediment.

2.3.  Ascension Island

St. Helena (British) (15° 58' S, 5° 43' W) lies slightly east of the mid-Atlantic Ridge, and is an isolated volcanic peak rising to a total of 5,223 m from the sea floor to its highest point at Diana Peak. With an area of 122 sq km it is a composite extinct volcanic cone, deeply eroded with radiating valleys intersected by deep valleys. The coastline ­comprises high stepped and sometimes vertical cliffs cut by steep v-sided valleys. The highest cliffs are 670 m at Flagstaff Bay in the north, where the valleys are marked by 200 m high coastal waterfalls, and at Man and Horse Cliffs in the south– west where they reach 580 m. The vertical steps in the coastal cliffs often correspond to differences in the composition of the outcropping lava flows. At Great Stone Top, Turks Cap and Prosperous Bay on the east coast, thick flows of trachyte tform vertical faces which ­cascade debris onto stone chutes below, some of which extend into the sea to form fringing small boulder beaches. Emerged shore platforms stand at 4–6 m above sea level and the numerous offshore islets and skerries common on the south and west coast, such as at Egg Island and Cockburn Battery, are almost absent elsewhere.

Ascension Island (British) (7° 57' S, 14° 22' W), lying in mid-ocean about 1,300 km north of Saint Helena, is the peak of a dormant Holocene stratovolcano that rises to a total of 3,860 m from the sea floor to its highest point at Green Mountain. More than 100 vents are found within its 98 sq km area and although eruptions continued into Quaternary times, there are no known historic eruptions. The climate is arid and vegetation is sparse, with scattered guano deposits. Imposing sea cliffs line the more exposed eastern coast, where Boatswain Bird islet, a monolithic stack of trachyte 98 m high, has a natural arch. The west coast comprises low lava fields, whose ­serrated seaward edge is truncated by low cliffs fronted by narrow and uneven shore platforms (>Fig. 6.5.3). However, Ascension also has superb white sand beaches derived from shell and coral. Little seems to be known of the nature of the sediment supply to the Ascension beaches, but the carbonates can only come from shells and coralline algae (such as

2.4.  St. Helena

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⊡⊡ Fig. 6.5.3 The west coast of Ascension Island is characterised by low gradient lava flows, whose termini have been eroded into low cliffs fronted by uneven platforms.

v­ egetation occurs above 1,350 m. The only township, Edinburgh, is located on a low-lying platform called Settlement Plain on the northwest side of the main island. Emerged beaches, platforms and caves are found at 5 m above sea level on Tristan and at 12 m above sea level on the small adjacent islands, indicating differential tectonic uplift. The cliffs are lower along the Edinburgh coast and a narrow gravel beach has developed along this stretch of coast, fed by material eroded from the cliffs. In the months following the 1961 eruption at Big Point, rapid recession of the cliff was noted (10 m in 8 weeks), this contributing to accretion of the downdrift gravel beaches which formed a spit enclosing a small lagoon. In some places, small and inaccessible beaches are found, some of which are sandy. Four small islands are associated with Tristan: Inaccessible (10 sq km; peak, 548 m), Nightingale (2.6 sq km; peak, 396 m), Stoltenhoff (0.25 sq km; peak, 105 m), and Middle (0.26 sq km; peak, 65 m), all are bordered by high lava and tuff cliffs with small and narrow sand and gravel beaches.

2.6.  Gough Island

Only in favoured locations, where the deep gullies reach the coast do small beaches, some of sand occur, such as at Lemon Valley Bay. At Sandy Cove on the south coast, the sandy beach is backed by calcareous sand dunes.

2.5.  Tristan da Cunha Tristan da Cunha (British) (37° 15' S, 12° 30' W) is a 98 sq km circular volcanic island, some 350 km NW of Gough Island. The base of the central volcanic peak lies 3,700 m deep on the sea-bed and the summit rises to almost 2,100 m above sea level at Queen Mary’s Peak. The lower slopes are almost entirely lava and form high cliffs that surround the island. Queen Mary’s Peak has about 30 parasitic scoria cones encircling the main crater with bordering cliffs up to 900 m high. The island is bleak and barren, but zoned

Gough Island (British) (40° 30' S, 10° 00' W), is also known as Diego Alvarez and lies 2,500 km east of the Cape of Good Hope. It is the eroded summit of a deeply dissected Tertiary volcano. The island is mountainous, rising to 910  m above sea level, and the lower slopes are covered with a luxuriant evergreen scrub forest that thins out above the 300-m contour. The steep slopes are cut by deep gullies or gulches descending to hanging valleys and waterfalls, and the coastline is characterised by steep cliffs that have been undercut to produce a variety of narrow boulder beaches at the foot of the cliffs. There are also lava arches, numerous islets, stacks (such as Tristania Rock (165 m)) and skerries that mostly lie within 100 m of the coast. Landings are prohibited and the islands bouldery beaches are ideal breeding grounds for fur seals.

2.7.  Bouvet Island Bouvet Island (Norway) (60 sq km; peak, 935 m) (54° 00' S, 3° 00' E) stands at the southern end of the mid-Atlantic ridge, and has a rugged ice-mantled Quaternary volcanic terrain with five glaciers, one of which ends on the east coast in ice cliffs up to 122 m high (>Fig. 6.5.4). The icecapped dome rises to 935 m. There is a central ice-filled crater, and the surface is more than 90% ice-covered. The island is extremely remote: there is no land in any direction for

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⊡⊡ Fig. 6.5.4 Bouvet Island. (Courtesy R. Muench, Earth and Space Research.)

more than 1,500 km. The adjacent seas have pack ice in winter, but are disturbed by frequent gales. Basalt cliffs cut back by stormy seas attain 490 m on the more exposed west coast, where in 1955–1957 an eruption or uplift added a new lava shelf 25 m above sea level from which a gravelly beach has been derived. Thompson Island, charted in the nineteenth century to the northeast, is thought to have disappeared as the result of a volcanic explosion.

3.  Islands of the Western Atlantic 3.1. South Shetland, South Orkney, South Sandwich and South Georgia Islands The Scotia Arc is a tectonically disrupted link between the fold mountain belts of South American cordillera and the  Antarctic Peninsula, and includes Late Palaeozoic– Mesozoic sedimentary formations with granitic intrusions and associated volcanics. The combined area of islands of the Scotia Arc is about 10,000 sq km, with peaks ­exceeding 2,000 m. South Shetland (4,700 sq km) is barren and heavily glacierised, the South Sandwich Islands (British) (310 sq km) have active volcanoes, and South Georgia (British) (3,756 sq km) is glacierised and mountainous, rising to 2,934 m at Mount Paget. South Georgia, South Orkney (British) and South Shetland Islands (British) contain

metamorphic gneisses, schists, and marble, but the South Sandwich islands are composed entirely of Cenozoic volcanics, and a submarine eruption north of Zavodovski Island in 1962 generated pumice deposits, which are found on the narrow beaches and were distributed far afield on the Antarctic Circumpolar Current. The islands are all barren and rugged, with glaciers, ice caps and Alpine peaks. Their coastlines are generally cliffed and rocky, but many inlets are found and plentiful glacial and glacifluvial debris provides the sediment sources for beach development. Rapid cliff retreat rates of 1 cm/year have been calculated for the cliffs of Fildes Peninsula in the South Shetland Islands, the result of efficient frost shattering. These coasts are affected by sea ice for up to 8 months of the year and the relative order associated with wave processes is annually swept away and replaced by the non-selective ­bulldozing and grinding action of floating ice, rendering beach landforms disorganised. In the South Shetland Islands, the coast is either rock or ice-cliff although some glacier termini rest on a low rocky platform at sea level. However, significant beach development occurs on the peninsulas that protrude beyond the ice cover and peninsulas on King George Island and Livingston Island are adorned with extensive emerged beach deposits, some of which connect offshore islets and skerries to the shore platforms of the main island. Prominent shore platforms are found at several altitudes including at

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present sea level, where extensive sub-horizontal shore platforms in sheltered locations give way to steeper, ramplike platforms in more exposed locations. This is a function of the persistence of floating ice blocks that horizontally abrade the surface in sheltered spots and wave activity that produces ramps in exposed sites. The upper part of the platforms is often adorned with an undulating layer of boulders that have been smoothed and compacted into extensive boulder pavements by the action of floating ice (Hansom and Gordon 1998). Some islands remain volcanically active and Penguin, Bridgeman and Deception are relatively new and active volcanoes. The sheltered inner crater of Deception Island has been inundated and has sand and gravel beaches in sharp contrast to the steep rocky outer coast. The coastline of the South Orkney Islands is formi­ dable and three of the four largest islands, Coronation, Powell and Laurie are dominated by large ice caps that spill over steep rocky cliffs, which plunge below sea level.  A  few bouldery beaches are found, for example, on Laurie Island. The remaining island, Signy, is low and rolling with a rocky sloping shore only with a few bouldery beaches. The South Sandwich group consists of a 400 km-long volcanic island arc that rises from a deep ocean trench. The 11 main islands are all volcanoes and the coastline is steep and cliffed with infrequent small bouldery beaches. Reefs and skerries abound especially between the smaller islands

in the extreme south at Thule, Cook and Bellingshausen islands. South Georgia lies on the northern part of the Scotia Arc, with several peaks rising to over 2,700 m. Long ridgelike and moraine-draped peninsulas are separated by deep glacial troughs (>Fig.  6.5.5), with the south coast more heavily glacierised than the north with steep cliffs and few small cliff-foot gravelly beaches. Glacier cover is more restricted on the north coast and the substantial ice-free peninsulas are flanked by high and steep cliffs, often with narrow shore platforms at their base together with numerous offshore islets and skerries. Several emerged marine platforms and beaches are found that relate to the glacial history but a systematic study of distribution and altitude is absent. The most extensive beaches of the Antarctic region lie on the north coast of South Georgia. Numerous glacifluvial streams deliver substantial amounts of material to extensive sand and gravel beaches, the largest of which are found at Salisbury Plain, Fortuna Bay, Cumberland Bay and St. Andrews Bay (Hansom and Gordon 1998).

3.2.  Falkland Islands The Falkland Islands (British) (52°S, 58°W) are also called Islas Malvinas and have a total area of 11,718 sq km. The coastline of the Falkland Islands, like the islands themselves, is low, flat and reminiscent of the coastline of the ⊡⊡ Fig. 6.5.5 The north coast of South Georgia has morainic plains fronted by gravelly beaches, in sharp contrast with the steep and heavily glacierised south coast.

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6.5

Outer Hebrides of Scotland. Although hundreds of smaller islands exist, the main island group comprises West and East Falkland separated by the steep-sided, fiord-like Falkland Sound (San Carlos Strait). Mt. Adam on West Falkland reaches 698 m. The intricate and crenulate nature of the lengthy coastline is related to the submergence and faulting of the underlying Devonian, Carboniferous and Permian limestones and sandstones. The landscape is treeless with extensive peat cover, and prominent periglacial stone runs of Pleistocene age are found (Hansom et al. 2008). In spite of the majority of the coastline being rocky, steep and high cliffs are mainly absent except in the west at Beaver Island where 100–200 m cliffs are found. In general, coasts are subdued, although the frequent westerly gales produce rapid recession where the rocky coast is replaced by periglacial sediment. Streams, although abundant and rarely dry, do not deliver much sediment to the beaches and the larger beaches are fed by shelly sand derived from nearshore shallow water. Significant beach development occurs on the outer coast and in many places strong winds have resulted in sand dune development (>Fig. 6.5.6). Many small sandy spits and tombolos connect small islets to the mainland.

Atol das Rocas (3° 50' S, 33° 40' W), (7.1 km2) lies 250 km off the Brazilian coast, but contrary to its name, it is not a coral atoll but an oval algal bank capping a seamount and enclosing a 6 m deep small lagoon. Two low sandy islands, Farol (106,000 sq m2) and Cemeteria (53,200 sq  m2), rise above the intertidal platform of algae, and have a capping of dune calcarenite. The highest point is a sand dune in the south of larger Farol Cay, at 6 m. Both islets are overgrown with grasses, bushes and palm trees.

3.3.  Trindade and Martin Vaz Rocks

3.5.  Fernando de Noronha

Trindade (8 sq km; peak, c. 600 m) (20° 30' S, 29°W) and the nearby Martin Vaz rocks are a complex of successively built volcanoes, the last (Paredo) a Holocene cone 217 m

The Fernando de Noronha archipelago (Brazilian) (3° 50' S, 32° 25' W) is volcanic and, with a main island (16.9 sq km) and about 20 associated islets, it is the emergent part

⊡⊡ Fig. 6.5.6 Sandy beaches on the east coast of East Falkland. Shallow water nearshore promotes beach development and strong winds carry the sand landward to produce dunes.

high on the southeast coast. The margins show high receding cliffs and truncated ravines, and a wave-cut tunnel 130 m long has been excavated in a volcanic headland. The exposed northeast coast has a coral reef, and there are beaches and dunes of calcareous sand as well as older dune calcarenites. Principe, on the south coast, has a sandy beach in front of a degraded cliff, and there is widespread evidence of about 3.5 m of emergence (platforms and beach rocks).

3.4.  Atol das Rocas

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of a large cone that rises 4,000 m from the sea floor. The relief is gentle, but the coastline is rocky and cliffed, with headlands in phonolite and bays carved in tuff. Wave action is strongest on the southeast coast, where Lith­ othamnion reefs have developed, while on the rugged north coast are beaches of sand, gravel and boulders, and some tombolos. Dune calcarenites and marine limestones are found above present sea level on the east coast. The tide range, generally small on oceanic islands, here attains 3.2 m at springs.

3.6.  Bermuda Bermuda (54 sq km) lies about 1,000 km east of Cape Hatteras on the east of the United States. It is a group of 150 limestone islands resting on the southern margin of a 650 sq km platform that is presently submerged to depths of 20 m. Beneath a thin capping of calcarenites, lava occurs at a variety of depths and the platform of volcanic rocks supports the northernmost coral reefs in the North Atlantic. Four geomorphological-ecological provinces are found: the 20 m depth reef-front terrace; the main 4 m deep coral-algal reef; the 16 m deep lagoon; and a northeast trending chain islands near the southern edge of the platform (Vacher 1973). The exposed limestones are mainly aeolian calcarenite, with only 10% beach and shallow water calcarenite. The supply of reef-derived sediment to beaches and dunes continues today but the present ­supply appears to be more restricted than in former times, since only on the island’s south shore do extensive beaches  backed by dune exist. These extensive sandy beaches are pink because of large numbers of fresh Homotrema clasts derived from the reefs, 1 km from the shoreline. Much of the south coast is cliff with only limited sediment. The sediment of the lagoon now plays little part in beach supply. This may be the reason why the lagoonfacing shoreline is erosional with a line of cliffs and few pocket beaches.

3.7.  Sable Island Sable Island (Canada) (43° 57' N, 60°W), lying on the shallow Sable Island Bank on the continental shelf 150 km southeast of Nova Scotia, is a low and wind-swept chain of sandy islands famous for their sandy shoals (>Fig. 6.5.7) and shipwrecks. The chain is 40 km by 1 km and of a variable but arcuate configuration. The Sable Island Bank is thought to be re-worked glacial sediment sourced from the maximum ice positions of the Late

Wisconsin and later readvances (King 2001). The islands have spectacular but desolate sandy beaches backed by sand dunes up to 30 m high (Byrne and McCann 1995). The dunes are primary in-situ dunes together with secondary dunes that have migrated across the island, a process exacerbated by a reduction in vegetation cover produced by the grazing of introduced ponies. Erosion during storms has caused eastward migration of 14.5 km over the past two centuries and necessitated the relocation of the lighthouse, although other buildings have been buried. However, deposition also occurs, and over the years, the land area has remained the same at about 30 sq km.

4.  Islands of the Eastern Atlantic 4.1.  Jan Mayen Dominated in the north by the 2,277 m stratovolcano, Beerenberg, the 54 km long-island of Jan Mayen (373 sq km) (Norway) is steep and rugged, comprising 200–400 m high cliffs eroded everywhere into ash, lava and tephra and in the north (Nord-Jan) into the ice of glaciers that flow from Beerenberg into the sea. The volcano remains active and the coastline has advanced as a result of lava tongues from  eruptions in 1970 and 1985, which have extended into the ocean. Such additions have resulted in pre-­existing  cliffs becoming buried by newer lavas that then become eroded by the sea, such as at Kroksletta, near the northern cape where the recent lava is now cliffed to a height of 5–13 m. Sud-Jan is a hilly cratered ridge of scoria and trachyte domes whose coast is composed of low but  steep rocky cliffs and low gradient but prominent rocky platforms connected by emerged and modern gravelly beaches. Unusually for a small mid-oceanic volcanic island, Nord-Jan is connected to Sud-Jan by a substantial 11 km-long barrier beach that encloses a lagoon on the south-central coast at Lagunevollen (>Fig. 6.5.8), with a smaller barrier and lagoon at Stasjonsbukta on the north coast. Although the north coast is subjected to variable amounts of sea ice for an average of 4 months, the south coast is mainly free of ice. Elsewhere, the rocky coastline of Sud-Jan is not as high as Nord-Jan and small gravelly cliff-foot beaches fed by wave erosion of the adjacent lavas are found.

4.2.  Faeroes The Faeroes (Denmark) (1,398 sq km) (62°N, 7°W) are 16 steep grassy volcanic islands whose landscape is the

Other Atlantic Ocean Islands

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⊡⊡ Fig. 6.5.7 Sable Island, Nova Scotia, is 40 km long and 1.5 km wide and is the exposed part of a sand shoal that extends for more than 160 km. (Courtesy National Aeronautics and Space Administration.)

⊡⊡ Fig. 6.5.8 South coast of Jan Mayen looking north towards the lower slopes of Beerenberg Volcano. (Courtesy Per-Einar Dahlen.)

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result of the erosion of past valley glaciers into a thick pile of horizontally-bedded Tertiary flood basalts. The horizontal nature of the lavas produces a stepped landscape whose interior is boggy. The islands are almost continuously exposed to storm waves, especially from the north and west and so the outer coast is erosional, rugged and dominated by high and plunging cliffs that reach 882 m at Slaettaratindur on Eysturoy. The rate of cliff recession is unknown, but is higher on the seaward facing sea cliffs than the immediately adjacent glaciated cliffs: Tindholmur island is effectively a mountain cut in half, with a vertical seaward cliff and a steep grassy slope on the landward side (>Fig.  6.5.9). Everywhere the coastline is indented with fiord-like inlets, deep clefts and geos and shore platforms are found intermittently. There are few beaches in the Faeroes and these are almost exclusively found at the sheltered fiord heads, where wave approach is unidirectional and small pocket beaches composed of black basalt sand, occasionally backed by low sand dunes are found, such as at Sandur and Saksun.

4.3.  Madeira Madeira (Portuguese) (32° 40' N, 17°W) consists of one large island and together with the smaller island of Porto Santo, the nearby Islas Desertas, and the uninhabited Selvagens to the south, they total 810 sq km and are the emergent summits of extinct (mainly Miocene) volcanoes

much dissected by deep ravines running down to the sea. Terraces mark stages in successive (Quaternary) uplift. The summit of Pico Rivo de Santana rises to 1,861 m. Madeiran slopes are steep, but well vegetated (annual rainfall, c. 600 mm, dry summer) and between the ravines they terminate in high coastal cliffs like Cabo Girao, west of Funchal, at 580 m among Europe’s highest (>Fig. 6.5.10). In the north, fragments of shore-platform are found as well as small islets and skerries, particularly near Faial. There are several well-developed stacks near Ponta do Sao Lourenco in the northeast. Small beaches of rounded grey basalt gravel are found at several places, particularly where river mouths occur. The only sandy beach is at Prainha, on the extreme east of the island where the sand is mainly basalt, but the south coast of nearby Porto Santo is dominated by a long white sand beach consisting of shelly material from the sea floor and aeolian sand blown in from the Sahara Desert. The Selvagens are composed of limestone capped by lava and ash and are cliffed to 100 m in the north, but with a gentler beach-fringed coast in the south. Emerged beaches and dune calcarenites are found on Pleistocene marine terraces at 3 and 7 m above sea level.

4.4.  Canary Islands The Canary Islands (Spanish), between 27° 30' and 29° 30' N and 13° 30'–18° 12' W, comprise seven main volcanic islands (total area, 7,500 sq km), six of which host volcanoes (Teneguia on La Palma, erupted in 1972). From east

⊡⊡ Fig. 6.5.9 The island of Tindholmur in the extreme west of Faeroe is what remains of an interfluve between two ice streams, the inner glacial trough retaining its shape while the outer one is an active marine cliff.

Other Atlantic Ocean Islands

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⊡⊡ Fig. 6.5.10 Cabo Girao, Madeira.

to west they are Lanzarote, Fuerteventura, Gran Canaria, Tenerife, La Gomera, La Palma, and Hierro. Pico de Teide on Tenerife (3,718 m) is the highest peak of these Atlantic islands but they all share a central volcanic spine falling steeply to a mainly rocky and cliff coast cut by deeply incised ravines. The main morphological features of the coastlines of the Canary Islands are related to vulcanicity: layered cliffs of superposed basalt and trachyte lavas, semicircular shapes corresponding to calderas, and long and low craggy volcanic peninsulas that have allowed beach development to occur. The islands have a dry subtropical climate, influenced by southeasterly trade winds from the Sahara: there is a little rain in November–December. Under the hot and dry northeast trade winds the eastern islands of Lanzarote and Fuertoventura receive dust and sand from the Sahara Desert 90 km away, resulting in white sand beaches rather than the black basaltic sand beaches of the other islands. The island slopes show radiating ravines that are short, steep, and gravelly. The north and west coasts are bold, with cliffs up to 300 m high in massive lavas, surrounded by fallen boulders. In the NW of Tenerife at Acantilado de Los Gigantes, the steep coastal slope rises to 500 m. Locally fans of gravelly material swept down ravines have been uplifted and are now cliffed, as on the north coast of Gran Canaria, where there are also emerged shore platforms. There are no biological reefs, but the north coast has ­sectors of shelly beach and calcareous dune sands.

Some of the dunes have been blown inland to form ­desertlike landscapes of bare, mobile dunes on Lanzarote and Fuerteventura. On the southeast coast valley mouth inlets are partly blocked by pebbly beaches, and there are incipient deltas; beaches of black volcanic sand are found locally. On the south coast of Gran Canaria there are beaches of shelly sand, which at Fataga enclose a freshwater lagoon. The tourist beach at San Andreas in SE Tenerife has been artificially nourished with sand shipped from the Sahara Desert. Pared is a sandy tombolo linking the former island of Jandia to mainland Fuertoventura.

4.5.  Cape Verde (Cabo Verde) Islands The Cape Verde Islands (14° 30'–17° 30' N and 22° 40'–25° 30' W), lie about 500 km west of Dakar in Senegal and have a very dry climate because of the northeast trade winds and the desiccating Harmattan winds from the Sahara, which are misty with airborne sand and dust. Mindelo on Sao Vicente has a mean temperature of 22°C in February and 27°C in September, and with a mean annual rainfall of 240 mm (mainly August/September), there are no permanent streams. Barren dissected volcanoes partly overlie Mesozoic marine limestone. There are 15 islands (total area, 4,033 sq km, coastline length 965 km), in two arcuate groups, the southern Sotavento group and the northern Barlavento group. Some of the islands in the

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east are low and undulating, such as Boa Vista, but those in the west are much higher and steeper, and surrounded by craggy plunging cliffs, such as Fogo. Strong winds mobilise sand (volcanic ash) to scour rock outcrops into rugged shapes. Some bays contain narrow fringing algal communities with patchy corals (gardens rather than reefs). Stormy seas have hurled boulders up to 200 m inland and undermined cliffs to produce massive rock falls. Emerged features have been noted at six levels up to 100 m. Four main stratovolcanoes are found in the Cape Verde Islands, Fogo, Brava, Santo Antao and San Vicente but only Fogo is active, its asymmetrical volcanic cone rising 2,829 m above the sea at Pico de Cano. The volcano has erupted ten times between 1500 and 1995. On its steeper east coast several lava flows formed in successive eruptions descend to basal cliffs, below which are small beaches of boulders and gravel. Porto Grande at Mindelo is a deep-water harbour in a volcanic crater, the edge of which has been breached and flooded by the sea. A similar picture characterises the other islands of the group, each being surrounded by steep cliffs cut by deep gullies at the foot of which occur infrequent gravelly beaches.

4.6.  Bioko (Fernando Po) The group of islands in the Gulf of Guinea (3° 30' N, 8° 40' W) include the large Bioko (Equatorial Guinea) (2,017 sq km), a high volcanic island densely covered in tropical rainforest, rising to a peak of 2,850 m at Santa Isabel, with steep rocky coasts interspersed with sheltered sandy bays and mangrove swamps. The capital, Malabo, stands by a bay formed where the sea has invaded a volcanic crater.

4.7.  Principe The 110 sq km island of Principe (Principe and Sao Tome) (1° 37' N,  7° 27' E) has a low, embayed north coast and a higher cliffed south coast. Pico de Principe in the south rises to  1,100 m. The island consists of Tertiary ­volcanic deposits, emergent since the Miocene with shelly ­limestone deposits at 130 m in the Forca Valley. It is dissected

by streams that deliver muddy waters to mangrove-fringed inlets behind barrier beaches, notably at Burras and Lapa.

4.8.  São Tomé Close to the Equator (0° 25' N, 6° 35' E), São Tomé (Principe and Sao Tome) is a large (854 sq km; peak, 2,023 m), densely vegetated island, with a generally subdued volcanic relief having occasional cones and plugs (Cao Grande, 663 m). Beaches of white calcareous sand interrupt high cliffs cut in basalt and phonolite, with columnar jointing well exposed in the Baio do Iologologo. To the south, Ilheu das Rolas is a much smaller volcanic island off the southern coast, and beyond is outlying Annó Bon, last of the group.

References Byrne ML, McCann SB (1995) The dunescapes of Sable Island. Canadian Landform Examples 31. Can Geogr 39:363–368 Guilcher A, Battistini R (1982) Erosional and constructional shore platforms and ancient beaches in the Azores. (in French). Revista Portuguesa de Geografia 15:221–242 Hansom JD, Gordon JE (1998) Antarctic environments and resources: a geographical perspective. Longman, London, 402p Hansom JD (2005) Atlantic Ocean islands, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Kluwer Academic Publications, Dordrecht, the Netherlands, pp 88–95 Hansom JD, Evans DJA, Sanderson DCW, Bingham R, Bentley M (2008) Constraining the age and formation of stone runs in the Falkland Islands using optically stimulated luminescence. Geomorphology 94:117–130 King RL (2001) A glacial origin for Sable Island. Geological Survey of Canada, Current Research D19:18 Mitchell-Thome RC (1970) Geology of the South Atlantic Islands. Beitrage zur Regionalen Geologie der Erde, 10. Gerbrüder Borntraeger, Berlin Mitchell-Thome RC (1976) Geology of the Middle Atlantic Islands. Beitrage zur Regionalen Geologie der Erde, 12. Gerbrüder Borntraeger, Berlin Smith HG, Hardy P, Leith IM, Spaull VW, Twelves EL (1974) A biological survey of St. Paul’s Rocks in the equatorial Atlantic Ocean. Biol J Linn Soc 6:89–96 Vacher L (1973) Coastal dunes of younger Bermuda. In: Coates DR (ed)  Coastal geomorphology. George Allen & Unwin, London, pp 355–391

7.0 British Isles – Editorial Introduction

The coasts of the British Isles are considered severally in the following chapters: England and Wales (> England and Wales Introduction, > Cumbria, > Lancashire, > Isle of Man, > North Wales, > West Wales, > South Wales, > Severn Estuary, > Gloucestershire, Somerset and North Devon, > North Cornwall, > Isles of Scilly, > South Cornwall, > South Devon, > Dorset, > Hampshire, > Isle of Wight, > Sussex, > Kent, > Essex, > Suffolk, > Norfolk, > Lincolnshire, > Yorkshire and Cleveland, > Durham, Tyne and Wear and > Northumberland ), > Scotland, > Northern Ireland, The Republic of > Ireland and the > Channel Islands.

1.  Coastal Geology A distinction was made by Mackinder (1902) between Highland and Lowland Britain, separated by a line drawn from the Exe estuary in South Devon to the Tees estuary on the North Sea coast. Highland Britain consists mainly of mountains and moorlands, dominated by Pre-Cambrian and Palaeozoic geological formations, Lowland Britain of escarpments and vales on Mesozoic, Tertiary and Quaternary formations. The outcrop of Permo-Triassic rocks marks a transition between the two (Stamp 1946). The coasts of Highland Britain (including Ireland and the Channel Islands) are steep and cliffy on the outcrops of older and harder rock formations, while the coasts of Lowland Britain are cliffed, where they intersect the sandstone, limestone and Chalk ridges and low-lying on intervening clay lowlands (Steers 1953). The landscape bears the imprint of Pleistocene glaciations, which at their maximum extended as far south as the Thames valley, South Wales and southern Ireland. Glacial drift deposits resulting from glaciation are widespread within this region. Where they reach the coast they are generally cliffed. In Highland Britain, there are cliffs where pockets of glacial drift occupy valleys, as on parts of the Welsh coast, notably in the northern part of Cardigan Bay, the Lleyn Peninsula, the Isle of Man, Lancashire and Cumbria, and locally in western Scotland. In Lowland Britain, cliffs cut in glacial drift are extensive in East Anglia and Holderness, and are found locally further north on the North Sea coast.

South of the limits of Pleistocene glaciation, there are rubble drifts (Head deposits) produced by intensive frost shattering. These periglacial deposits mantle hillsides and coastal slopes, and are well seen in SW England, where they have been undercut to a varying extent (related to exposure to strong wave action), forming slope-over-wall coast profiles. Similar coastal slopes existed in south-east England, but have been generally destroyed by Holocene marine erosion, and persist only where they have been protected by Holocene deposition, as in the lee of Slapton Sands in South Devon and Chesil Beach in Dorset, and behind alluvial lowlands such as Romney Marsh. Slopeover-wall profiles are also found on parts of the coasts of Highland Britain, as in Cardigan Bay, where periglaciation followed glacial deposition.

2.  Coastal Environments With a prevalence of westerly winds, Highland Britain has higher precipitation (rain and snow) than Lowland Britain, which is comparatively dry. Tide ranges are macrotidal (4–6 m) and megatidal (>6  m) around much of the British coastline, reducing to mesotidal (2–4 m) in NW Wales and W Scotland, Orkney and Shetland and NE Scotland and E Norfolk, and microtidal ( Cumbria to > Cornwall), west to east (> Cornwall to > Kent) and south to north (> Essex to > Northumberland). In England and Wales most of the stages in the geological column are somewhere represented. A notable gap is the Miocene (5–23 million years ago), when Britain was a land area, subject to erosion rather than deposition, apart from some gravel deposits, while some stages, such as the Rhaetic and the Palaeocene, are better developed elsewhere, particularly in Europe. There has been tectonic uplift in the north and west of Britain and subsidence in the south and east. This transverse tilting results in the oldest (Pre-Cambrian) rocks outcropping in parts of Anglesey while the youngest are in East Anglia. The pattern has been complicated by three phases of tectonic activity, the Caledonian, producing structures that run SW-NE in Wales and Cumbria, the Hercynian, producing east-west structures in South Wales and South-West England, and the Alpine, also producing east-west structures across southern England, including the Weald and the Hampshire Basin. The result is that the older rock formations (PreCambrian and Palaeozoic) form the uplands of Wales and the south-west peninsula, and the overlying Mesozoic strata dip generally south-east, the more resistant formations outcropping in roughly parallel escarpments between vales cut in weaker formations in a sequence of diminishing age up to the Chalk. The Tertiary formations are largely confined to the synclinal London and Hampshire Basins in the transversely folded Chalk. The various structures have been truncated by erosion to form the present coastline, and thus expose many different geological formations in cliffs and rocky shores. The influence of Caledonian structures is shown by the elongated SW-NE rock outcrops in Anglesey and the Lleyn

Peninsula in North Wales. The Hercynian structures of the Mendip Hills produce the ridges that protrude into the Bristol Channel at Weston super Mare. The Alpine folding is responsible for the high, narrow Chalk ridge that runs through the Isle of Wight, protruding as cliffy headlands at the eastern and western extremities: Culver Cliff and The Needles. Cambrian rocks outcrop in the Isle of Man, the Lake District and Snowdonia, and Ordovician and Silurian rocks over much of Wales. Devonian and Carboniferous rocks are exposed in Devon and Cornwall, South Wales, and Northumbria. These various formations can be seen in the cliffs and rocky shores of the western coasts of England and Wales. The Permian and Triassic outcrops, which form a low zone between Upland Britain to the west and Lowland Britain to the east, outcrop on the Devon coast between Torquay and Seaton and on the north-east coast between the mouth of the River Tees and South Shields. The Jurassic formations are well exposed in the cliffs of Lyme Bay between Seaton and Portland Bill, in Weymouth Bay and Ringstead Bay, and locally to the east, particularly between Lulworth Cove and Swanage. They are also seen on the north-east coast between Filey Bay and Redcar. The Lower Cretaceous formations have been brought up above sea level in upfolded areas in Purbeck, the Isle of Wight and The Central Weald. In Dorset they outcrop on either side of Lulworth Cove, in Worbarrow Bay and in Swanage Bay, and in Sussex and Kent they extend from Eastbourne to Folkestone. Above them the Chalk, the most distinctive of the Cretaceous formations, outcrops in sectors of cliffed coast at Beer Head in Devon, between White Nothe and Worbarrow Bay in Dorset. The Chalk passes below sea level in the Hampshire Basin, but rises sharply to form a central ridge between The Needles and Culver Cliff in the Isle of Wight. The South Downs and North Downs are in-facing Chalk escarpments overlooking the Weald and reaching the Sussex coast between Brighton and Eastbourne and the Kent coast from west of Dover round to Deal. Chalk cliffs reappear in Thanet extending from Pegwell Bay past Ramsgate and Margate to Minnis

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Bay, west of Birchington. On the North Norfolk coast there is ice-rafted Chalk in the cliffs between East and West Runton, and Chalk outcrops on the shore platform at Sheringham. In Yorkshire the Chalk of the Wolds comes to the coast at Flamborough Head, with Chalk cliffs extending from Sewerby (NE of Bridlington) round to Bempton. The Tertiary formations (Palaeocene, Eocene and Oligocene) overlie the Chalk, and occur in the Hampshire Basin and the Isle of Wight, where they outcrop in cliffs along the northern coast. They are also seen on the north coast of the Isle of Thanet, and in East Anglia, where the cliffs near Walton-on-the-Naze are cut in Eocene formations, while Pliocene formations outcrop in cliffs north of Aldeburgh in Suffolk and on the Norfolk coast between Trimingham, Cromer, Sheringham and Weybourn. Pleis­­ tocene formations are dominated by glacial drift, which has been cliffed on several parts of the east coast of Britain, notably in East Anglia and Holderness, and locally on the west coast bordering the Irish Sea and on the northern peninsula of the Isle of Man. There are also Pleistocene periglacial drift deposits, notably the Head formations prominent on coastal slopes in south-west England and South Wales.

2.  The Coastline The combined coastline of England and Wales (>Fig. 7.1.1) is about 2,734 miles (4,400 km) in length. Its outlines are related to geological outcrops, the more resistant rocks forming headlands while weaker formations have been excavated as bays and inlets (Clayton and Shamoon 1998). Palaeozoic rocks dominate the intricately dissected western peninsulas of Cumbria, Wales and S.W. England, and a sequence of Mesozoic sandstones and limestone cuestas with intervening clay vales ending in alternating cliffs and lowlands on the southern and eastern coasts. Tertiary sedimentary formations outcrop only locally in Dorset, Hampshire and Thanet, but Pleistocene glacial deposits are extensive on the east coast, especially in Holderness and East Anglia. A general account was provided by Steers (1964), and a review of scientific sites by May and Hansom (2003). The cliffs and depositional features formed during the later stages of the Flandrian (Late Pleistocene to Holocene) marine transgression and its aftermath (during the past 5,000 years) of relative stillstand, but relict Pleistocene raised beaches and periglaciated slopes persist where erosion has been slow, mainly on the western coast. Remains of

forests submerged by the Flandrian marine transgression are seen at low tide in the Bristol Channel and ­elsewhere, and continuing subsidence in south-eastern England is indicated by the foundations of Roman London, now 2–3 m below high-tide level. Tide gauge records show that northern England has been rising and the south and south-east sinking at rates of a few millimetres per year. There is much debate on the significance of actual sea level oscillations, of tectonic uplift and subsidence, of upward and downward isostatic movements in response to deglaciation and rising sea levels, of erosion and sedimentation, of land reclamation and coastal engineering as factors influencing ongoing changes in the relative levels of land and sea around England and Wales. The rivers of England and Wales flow into estuaries formed by submergence of valleys that had been deepened during Pleistocene low sea level phases, and fluvial deposition has formed valley floors, but there are no protruding deltas. Embanking and reclamation of intertidal areas has been extensive, especially in Romney Marsh north of Dungeness, the Somerset Levels and the Fens south of The Wash. Marine sand has been washed into several of the drowned valley mouths (rias) of south-west England, producing broad intertidal sandflats, but those that open on sheltered parts of the coast, such as Carrick Roads in Cornwall and Milford Haven in Pembrokeshire, have escaped this accretion, and remain as deep inlets. Wave energy is strong from the southwest through the Atlantic approaches, diminishing into the English Channel and through St. George’s Channel to the Irish Sea. The east coast is sheltered from the prevailing westerly winds, but receives waves from northerly and easterly directions over the North Sea. Refracted patterns of dominant waves have shaped curved bays and beaches, as in Lyme Bay on the Dorset coast (Bird 1989). Longshore drifting of beach material is eastward along the south coast of England and southward down the east coast, except in north Norfolk where westward movement of shingle has contributed to the intermittent growth of compound recurved formations at Blakeney Point and Scolt Head Island. Beach drifting and cliff recession are most pronounced when onshore winds and low barometric pressure bring large waves in at unusually high levels. Major storm surges have occurred on the east coast several times in the past few centuries, the most recent being in 1953 and 1978. The duration and intensity of wave action on the shore is also influenced by tidal regimes. Mean spring tide ranges (>Fig. 7.1.1) are generally between 3 and 5 m, augmented in the Bristol Channel to more than 12 m at Avonmouth, with tidal bores transmitted up the Severn, Avon, and

England and Wales

⊡⊡ Fig. 7.1.1 The coastline of England and Wales. (Courtesy Geostudies.)

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Parrett rivers, and falling to less than 2 m on the central south coast, where double high tides occur. Strong currents are generated in the Irish Sea and through Menai Strait, and tidal currents have shaped complex shoal and channel topography off the east coast, where wave action is relatively weak. Tidal variations also influence the ecological zonations of plant and animal communities on rocky shores (Lewis 1964). Cliffs are found on a variety of geological formations where uplands have been truncated by marine erosion. Much of the coastline of England and Wales runs across the geological strike, so that many formations outcrop in rapid succession along the coast. The most characteristic are the vertical white cliffs of thick gently dipping Cretaceous Chalk on either side of the transected Weald, between Brighton and Eastbourne and between Dover and Deal; however, cliffs in strongly folded chalk occur on the Dorset coast and the Isle of Wight, while at Flamborough Head chalk cliffs are capped by slopes in glacial drift, a sequence that results in slope-over-wall profiles. Steep and high cliffs are also well developed on limestones at either end of the Jurassic outcrop that crosses England from Dorset to Yorkshire, with an outlier in Glamorgan; on the red Permian sediment of St. Bees Head and East Devonshire; and on the Tertiary sands of Bournemouth Bay and the Isle of Wight. Sloping cliffs (Fig.  7.1.2) were reviewed by a Royal Commission in 1911, which concluded that over a period of about 35 years the gains (14,344 ha) largely by reclamation around estuaries substantially exceeded the losses by erosion (1,899 ha). Historical evidence shows that on the outer coastline only a few sectors (spits, cuspate forelands, accretion alongside breakwaters) have prograded, while on the Chalk (May 1971) and on the softer formations along the  east and south coast rapid cliff recession has consumed  much agricultural land, several settlements (e.g. the fourteenth-­century port of Dunwich, Suffolk) having been lost (Bird and May 1976). It is difficult to find beaches that are actively prograding, and the seaward margins of coastal dunes are commonly cliffed by erosion during recent decades. The response to coastal erosion has been extensive construction of sea walls and boulder ramparts, to­­gether with groynes intended to retain protective beaches, but sea walls and boulder ramparts have often led to beach depletion and shore lowering by wave reflection scour while (as has been noted) groynes impede longshore drifting and result in erosion down-drift. In recent years, coastal management has begun to take account of coastal sediment cells, within which longshore, onshore and offshore movements of sand and gravel occur, and can be assessed in terms of sediment budgets. On this basis, sand and gravel from offshore or hinterland sources can be deposited to augment beaches, and reference has been

7.2

made to such artificial beach nourishment on parts of the coast as an alternative method of coast protection. Predicted changes in climate and sea level are likely to lead to substantial changes on the coast of England and Wales (Cracknell 2003). Low-lying areas, including many that have been embanked and reclaimed during the past few centuries, are likely to be submerged by the sea, and erosion will develop and intensify on cliffy coasts, beaches and dunes, and marshlands. One response may be to extend and elaborate coastal defences, building up lowlying areas by infilling, but another may be managed retreat, by which coastal areas of little economic value will be abandoned to marine incursion.

References Bird ECF (1989) The beaches of Lyme Bay. Proc Dorset Nat Hist Archaeol Soc 111:91–97 Bird ECF, May VJ (1976) Shoreline changes in the British Isles during the past century. Bournemouth College of Technology, Bournemouth, England Clayton KM, Shamoon N (1998) A new approach to the relief of Great Britain, 1. The machine-readable database. Geomorphology 25:31–42 Cracknell BE (2003) Outrageous waves: global warming and coastal change in Britain. Phillimore, Chichester, West Sussex Gresswell RD (1953) Sandy shores in South Lancashire. University of Liverpool, Liverpool King CAM (1972) Beaches and Coasts. St. Martin’s, London Lewis JR (1964) The ecology of rocky shores. English Universities Press, London May VJ (1971) The retreat of chalk cliffs. Geogr J 137:203–206 May VJ, Hansom JD (2003) Coastal geomorphology of Great Britain. Geological Conservation Review Series 28. Joint Nature Conservation Committee, Peterborough, ON, pp 737 Steers JA (ed) (1960) Scolt Head Island. Cambridge University Press, Cambridge Steers JA (1964) The coastline of England and Wales. Cambridge Uni­ versity Press, Cambridge Wright LW (1970) Some characteristics of the shore platforms of the English Channel. Zeitschrift für Geomorphologie 11:36–46

7.2  Cumbria 1.  Introduction The mountains of the Lake District have developed on Lower Palaeozoic slates and volcanic rocks, and have been shaped by Pleistocene glaciation into peaks and ridg­ ­es, separated by the deep troughs which contain the lakes. The Lower Palaeozoic formations have had little influ­­ ence  on the Cumbrian coastline. On the northern side, they pass beneath Carboniferous Limestone and the Coal

Measures, which reach the coast between Maryport and Whitehaven, and on the southern side, they pass beneath Carboniferous Limestone, which forms the hills of the Furness Peninsula, Grange-over-Sands and the Arnside district. There is a fringe of Triassic sandstones and marls which reaches the north coast near Maryport, and fringes Solway Firth is prominent on St Bees Head and underpins the coast down to Barrow-in-Furness and the Isle of Walney, with an outlier on the Flookburgh Peninsula.

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hybridization took place in Southampton Water between local Spartina maritima and American Spartina alterniflora, yielding a vigorous new species, Spartina anglica. This spread rapidly and was introduced to Poole Harbour and other estuaries in England and Wales, and became established on board intertidal mudflats in Bridgwater Bay. It has accelerated the building of depositional terraces, but in some areas it is dying back and the marshlands are eroding. Changes in the outline of England and Wales (>Fig.  7.1.2) were reviewed by a Royal Commission in 1911, which concluded that over a period of about 35 years the gains (14,344 ha) largely by reclamation around estuaries substantially exceeded the losses by erosion (1,899 ha). Historical evidence shows that on the outer coastline only a few sectors (spits, cuspate forelands, accretion alongside breakwaters) have prograded, while on the Chalk (May 1971) and on the softer formations along the  east and south coast rapid cliff recession has consumed  much agricultural land, several settlements (e.g. the fourteenth-­century port of Dunwich, Suffolk) having been lost (Bird and May 1976). It is difficult to find beaches that are actively prograding, and the seaward margins of coastal dunes are commonly cliffed by erosion during recent decades. The response to coastal erosion has been extensive construction of sea walls and boulder ramparts, to­­gether with groynes intended to retain protective beaches, but sea walls and boulder ramparts have often led to beach depletion and shore lowering by wave reflection scour while (as has been noted) groynes impede longshore drifting and result in erosion down-drift. In recent years, coastal management has begun to take account of coastal sediment cells, within which longshore, onshore and offshore movements of sand and gravel occur, and can be assessed in terms of sediment budgets. On this basis, sand and gravel from offshore or hinterland sources can be deposited to augment beaches, and reference has been

7.2

made to such artificial beach nourishment on parts of the coast as an alternative method of coast protection. Predicted changes in climate and sea level are likely to lead to substantial changes on the coast of England and Wales (Cracknell 2003). Low-lying areas, including many that have been embanked and reclaimed during the past few centuries, are likely to be submerged by the sea, and erosion will develop and intensify on cliffy coasts, beaches and dunes, and marshlands. One response may be to extend and elaborate coastal defences, building up lowlying areas by infilling, but another may be managed retreat, by which coastal areas of little economic value will be abandoned to marine incursion.

References Bird ECF (1989) The beaches of Lyme Bay. Proc Dorset Nat Hist Archaeol Soc 111:91–97 Bird ECF, May VJ (1976) Shoreline changes in the British Isles during the past century. Bournemouth College of Technology, Bournemouth, England Clayton KM, Shamoon N (1998) A new approach to the relief of Great Britain, 1. The machine-readable database. Geomorphology 25:31–42 Cracknell BE (2003) Outrageous waves: global warming and coastal change in Britain. Phillimore, Chichester, West Sussex Gresswell RD (1953) Sandy shores in South Lancashire. University of Liverpool, Liverpool King CAM (1972) Beaches and Coasts. St. Martin’s, London Lewis JR (1964) The ecology of rocky shores. English Universities Press, London May VJ (1971) The retreat of chalk cliffs. Geogr J 137:203–206 May VJ, Hansom JD (2003) Coastal geomorphology of Great Britain. Geological Conservation Review Series 28. Joint Nature Conservation Committee, Peterborough, ON, pp 737 Steers JA (ed) (1960) Scolt Head Island. Cambridge University Press, Cambridge Steers JA (1964) The coastline of England and Wales. Cambridge Uni­ versity Press, Cambridge Wright LW (1970) Some characteristics of the shore platforms of the English Channel. Zeitschrift für Geomorphologie 11:36–46

7.2  Cumbria 1.  Introduction The mountains of the Lake District have developed on Lower Palaeozoic slates and volcanic rocks, and have been shaped by Pleistocene glaciation into peaks and ridg­ ­es, separated by the deep troughs which contain the lakes. The Lower Palaeozoic formations have had little influ­­ ence  on the Cumbrian coastline. On the northern side, they pass beneath Carboniferous Limestone and the Coal

Measures, which reach the coast between Maryport and Whitehaven, and on the southern side, they pass beneath Carboniferous Limestone, which forms the hills of the Furness Peninsula, Grange-over-Sands and the Arnside district. There is a fringe of Triassic sandstones and marls which reaches the north coast near Maryport, and fringes Solway Firth is prominent on St Bees Head and underpins the coast down to Barrow-in-Furness and the Isle of Walney, with an outlier on the Flookburgh Peninsula.

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Glacial drift deposits are extensive on the lowlands around the Lake District, and prominent in low cliffs along parts of the coastline, as on the Isle of Walney (Huddart et al. 1977). Mean spring tide range at Glasson on Solway Firth is 5.2 m, increasing to 8.3 m at Silloth, then declining to 7.7 m at Maryport, 7.5 m at Workington and 7.3 m at Whitehaven. Along the coast south past St Bees the range remains around 7 m, with 7.5 m at Tarn Point and 7.6 m on the Duddon Bar, while the port of Barrow in Furness has 8.2 m.

2.  The Coastline The Cumbrian coast begins at the head of Solway Firth, where the rivers Sark, Esk and Eden flow to a coast with wide salt marshes declining seaward to mudflats exposed at low tide. The Solway salt marshes, like others on the west coast of Britain, are on firm sand or muddy sand, and form lawn-like swards when grazed. Solway Firth widens ­rapidly west­ward, and salt marshes and tidal mudflats are  extensive. There are successive salt marsh terraces (>Fig.  7.2.1), rising landward and related to isostatic uplift, complicated by the effects of migrating river channels and a recent slow rise in relative sea level. The seaward margins of these salt marshes are often cliffed and eroding, despite the emergence. There are scattered scars (boulders washed out of the glacial drift) along the shore. Moricambe Bay has

extensive sandy flats backed by high salt marsh and peat bogs. Numerous small parallel tidal creeks drain the salt marsh, which has cliffed or steep seaward margins exposing layered mud. There are a few Spartina clones on the mudflats. Grune Point is a dune-capped spit of sand and shingle that constricts the entrance to Moricambe Bay. It has been supplied by longshore drifting NE from Silloth, indicated by angled beach segments between groynes. SW from Silloth the sandy intertidal zone widens to an area of ridge and runnel exposed at low tide. The coast is bordered by dunes underlain by glacial drift, from which the shingle and sand beaches, nearshore sand bars and boulder scars have been derived. A 7.6-m emerged beach south of Dubmill Point, has Holocene sand, clay and peat deposits, thought to have formed in a dune-fringed lagoon during the Late Qua­ ternary (Flandrian) marine transgression, and been subsequently raised by postglacial isostatic rebound. In general, emerged beaches decline in level southward along the Cumbrian coast. Allonby Bay has low cliffs cut in boulder clay with occasional outcrops of Triassic St Bees Sandstone. At Maryport, a transverse fault brings up Carboniferous rocks, Millstone Grit and Coal Measures. Near Workington, there are segments of cliff cut into basic slag cliff dumped from a former steelworks at Moss Bay, which have yielded grey sand and gravel and dark brown ironstone to adjacent beaches (>Fig. 7.2.2). South of Workington, the coast was modified by coal mining. The Coal Measures dip beneath sandstones at ⊡⊡ Fig. 7.2.1 A grassy terrace, microcliff and salt marsh at Cardurnock, Solway Firth. (Courtesy Geostudies.)

Cumbria

7.2

⊡⊡ Fig. 7.2.2 Slag heaps at Moss Bay, Workington. (Courtesy Geostudies.)

Whitehaven, and to the south Permian and Triassic sandstones form the cliffed St Bees peninsula, rising to 98 m on North Head (Eastwood et al. 1931). Cliffs in stratified red  sandstone decline to Flenswick Bay at the incised mouth of a small valley. There is a grey pebble beach and the cliffs  have basal ledges of seaward-­dipping Triassic sandstone that show differential erosion along joint planes (>Fig.  7.2.3). The cliffs are locally ­overhanging, and an abrasion notch is prominent along their base. The cliffs then rise towards South Head (>Fig. 7.2.4), where the St Bees Sandstone dips landward and forms an upper cliff, then decline towards the valley of Rottington Beck, which contains glacial drift, and to the south cliffs and bluffs in boulder clay extend for several kilometres behind a grey shingle beach. Three rivers, the Irt, Mite and  Esk, converge into the estuary at Ravenglass, the Irt ­de­­flected southward by the dune-capped spit at Drigg Point. At low tide, the estuary is bordered by gravels  and  sand with mud below mid-tide level. Eskmeals Point to the south, is also dune-covered, and Selker Point is a low eroding cliffed promontory of glacial drift (>Fig. 7.2.5). The River Duddon opens into a large funnel-shaped estuary where at low tide the Duddon Sands are exposed up to 3 km wide and traversed at low tide by the narrow winding river channel and some smaller tidal creeks. On the east coast of the Duddon estuary, salt marshes are backed by steep slopes in Carboniferous Limestone. The marshes pass southward behind a sand dune fringe at Sandscale Haws and Lowsy Point is a dune-covered spit

bordering the shallow northern entrance to the channel behind the Isle of Walney. The Isle of Walney has the form of a barrier island, but it consists of glacial drift. The western part of the island is a series of low hills of glacial drift linked by barriers of sand and gravel, and the northern end is a sandy spit capped by dunes, North End Haws. The Isle of Walney has been frequently overwashed by the sea and coastal erosion has been severe. Parts of the coast have been protected against erosion by sea walls and concrete blocks. The southern part is narrower, with low cliffs that fade as the coast swings towards South End Haws. To the east of the island, an extensive area of intertidal sand and ­mudflats is split by the Walney Channel, which gives ships high tide access to the industrial port of Barrowin-Furness. Morecambe Bay extends across to Heysham in ­Lancashire. The coast runs northward past Piel Island and Roa Island with segments of cliff cut in glacial drift (including truncated drumlins, as at Wadhead Hill), ­fringing salty  mar­­shes and bouldery scars. Beyond Ulverston is ­Greeno­­dd, an old port below the confluence of the Rivers Crake and Leven, which drain from Coniston Water and Lake Windermere respectively, and open into  a  wide estuary with channels and shoals bordered by  salt marsh. Notices warn of quicksands and rapid incoming tides. The hilly country to the east of the estuary is on Carboniferous Limestone, but the undulating  country to the south, around Flookburgh, includes drumlins.

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⊡⊡ Fig. 7.2.3 Flenswick Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.2.4 Shingle beach and cliffs in dipping St Bees Sandstone south of Flenswick Bay, looking towards South Head. (Courtesy Geostudies.)

Humphrey Point stands at the southern end of a prominent Carboniferous Limestone ridge, Humphrey Head, rising to a summit of 53 m. The limestone dips eastward, so that the western side is a high escarpment cliff (>Fig. 7.2.6), while the eastern side is a gentler dip-slope, declining to a  low basal ramp with low wave energy because of the wide salt marsh. Humphrey Point projects

as a rocky spur south-ward into the mudflats of Morecambe Bay (>Fig. 7.2.7). The coastline turns northward, bordered by salt marshes, past Grange-over-Sands. Where there are views across Morecambe Bay, with vast sand flats exposed at low tide. It has tidal channels that change rapidly, deepening, shallowing and migrating. There are sandflats firm enough

Cumbria

⊡⊡ Fig. 7.2.5 Cliff recession at Selker Point has cut into a house. (Courtesy Geostudies.)

⊡⊡ Fig. 7.2.6 Escarpment cliff of Carboniferous Limestone on the western side of Humphrey Head. (Courtesy Geostudies.)

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Lancashire, Merseyside and Cheshire

⊡⊡ Fig. 7.2.7 Spur of Carboniferous Limestone at Humphrey Point, projecting into the tidal flats of Morecambe Bay. (Courtesy Geostudies.)

to drive a tractor across, but there are also areas of soft mud, particularly where fresh water flows out from rivers. East of the Kent estuary, a limestone ridge steep wood­ ­ed slopes descending to Arnside Point. The Cumbrian coast ends where the Lancashire boundary crosses the Silverdale marshes.

References Eastwood T, Dixon EEL, Hollingworth SE, Smith D (1931) The geology of the Whitehaven and Workington district. Memoirs of the Geological Survey, London Huddart D, Tooley MJ, Carter PA (1977) The coasts of northwest England. In: Kidson C, Tooley MJ (eds) The Quaternary history of the Irish sea. Seel House Press, Liverpool, pp 119–154

7.3  Lancashire, Merseyside and Cheshire 1.  Introdution The Lancashire, Merseyside and Cheshire coast is generally low-lying and dominated by glacial drift deposits, with only local outcrops of the underlying solid formations. It is notable for its wide sandy estuaries, Morecambe Bay, the Lune, the Ribble, the Mersey and the Dee, at the mouths of valleys, which were incised during Pleistocene low sea level phases, submerged by the Late Quaternary (Flandrian) marine transgression, and then partly filled with sediment brought down by the rivers or swept in from the sea floor. Shore and near shore sediment range from shingle and sand to silt and clay, derived from the glacial drift by wave action as the tides (which are substantial) rise and fall. Salt marshes are extensive in the upper intertidal zone, and there has been much embanking and reclamation of these

to secure for farmland or develop ports and industrial areas. Mean spring tide range is generally large, 8.4 m at Morecambe, Heysham and Fleetwood, 7.5 m at St. Anne’s, 8.3 m at Liverpool and 7.8 m at Hilbre Island. In the Dee estuary, the mean spring tide range at Connah’s Quay is 4.6 m.

2.  Lancashire The Lancashire coast begins near the northern end of Silverdale Marsh, which is backed by steep wooded bluffs of Carboniferous Limestone marking a former coastline that extends round to Jenny Brown’s Point. At Silverdale the salt marsh fringe is eroding, and a gravelly beach is forming in front of the eroded cliff (>Fig. 7.3.1).

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Lancashire, Merseyside and Cheshire

⊡⊡ Fig. 7.2.7 Spur of Carboniferous Limestone at Humphrey Point, projecting into the tidal flats of Morecambe Bay. (Courtesy Geostudies.)

to drive a tractor across, but there are also areas of soft mud, particularly where fresh water flows out from rivers. East of the Kent estuary, a limestone ridge steep wood­ ­ed slopes descending to Arnside Point. The Cumbrian coast ends where the Lancashire boundary crosses the Silverdale marshes.

References Eastwood T, Dixon EEL, Hollingworth SE, Smith D (1931) The geology of the Whitehaven and Workington district. Memoirs of the Geological Survey, London Huddart D, Tooley MJ, Carter PA (1977) The coasts of northwest England. In: Kidson C, Tooley MJ (eds) The Quaternary history of the Irish sea. Seel House Press, Liverpool, pp 119–154

7.3  Lancashire, Merseyside and Cheshire 1.  Introdution The Lancashire, Merseyside and Cheshire coast is generally low-lying and dominated by glacial drift deposits, with only local outcrops of the underlying solid formations. It is notable for its wide sandy estuaries, Morecambe Bay, the Lune, the Ribble, the Mersey and the Dee, at the mouths of valleys, which were incised during Pleistocene low sea level phases, submerged by the Late Quaternary (Flandrian) marine transgression, and then partly filled with sediment brought down by the rivers or swept in from the sea floor. Shore and near shore sediment range from shingle and sand to silt and clay, derived from the glacial drift by wave action as the tides (which are substantial) rise and fall. Salt marshes are extensive in the upper intertidal zone, and there has been much embanking and reclamation of these

to secure for farmland or develop ports and industrial areas. Mean spring tide range is generally large, 8.4 m at Morecambe, Heysham and Fleetwood, 7.5 m at St. Anne’s, 8.3 m at Liverpool and 7.8 m at Hilbre Island. In the Dee estuary, the mean spring tide range at Connah’s Quay is 4.6 m.

2.  Lancashire The Lancashire coast begins near the northern end of Silverdale Marsh, which is backed by steep wooded bluffs of Carboniferous Limestone marking a former coastline that extends round to Jenny Brown’s Point. At Silverdale the salt marsh fringe is eroding, and a gravelly beach is forming in front of the eroded cliff (>Fig. 7.3.1).

Lancashire, Merseyside and Cheshire

7.3

⊡⊡ Fig. 7.3.1 The eroding salt marsh at Silverdsle. (Courtesy Geostudies.)

The salt marshes extend along the NE shore of Morecambe Bay, across the Kent estuary and round to Scalestones Point. The mean spring tide range here is over 8 m and the tide rises and falls rapidly across muddy sands that are up to 10 km wide at the lowest tides. The salt marshes have been partially drained and embanked, and are grazed by cattle. Nevertheless, the salt marsh fringe has been eroding. South of Scalestones Point, the eroded coastline has been heavily armoured with large limestone block ramparts, with groynes intended to retain a beach. The beach remains sparse and gravelly, with extensive muddy sand exposed at low tide. At Morecambe, attempts have been made to improve it by dumping sand and shingle in front of the promenade. South of Morecambe, there are low cliffs and a rocky shore in Millstone Grit at Heysham. To the south, the coast is low and marshy with many tidal creeks, and towards Sunderland Point, there is a shingle beach. It becomes a spit on the northern side of the Lune estuary, which is bordered by salt marshes (>Fig. 7.3.2). The River Cocker opens through the marshes to a little gulf at Bank End, and at low tide, the wide Cockerham Sands are exposed. A long embankment begins on the south bank of the River Cocker, enclosing reclaimed land. It is armoured with large limestone boulders along a 3 mile stretch westward from Fluke Hall, where the salt marsh fringe fades out. A concrete sea wall then extends

along the Esplanade at Knott End on Sea, which at low tide is fronted by rough grassy salt marsh. This grades into a wide sandy area known as Preesall Sands and Bernard Wharf, besides the mouth of the River Wyre. West of the River Wyre, the coastline is artificial, with a  high embankment, concrete sea walls and numerous groynes extending past Fleetwood to Cleveleys and Black­ pool. It borders a peninsula, which is of glacial drift underlain by Keuper Marl. The shingle fades out towards Blackpool, where the famous Blackpool Sands are largely submerged at high tide. To the south of Blackpool, the coast swings eastward past Lytham St. Annes, where it passes into salt marshes on the northern shore of the Ribble estuary.

3.  Merseyside The southern shore of the Ribble estuary has salt marshes behind wide intertidal Marshside Sands. The coast be­­ comes sandy towards Southport, where the sands are calcareous, incorporating the residues of shelly marine organisms as well as inorganic sand derived from glacial drift deposits. The coastline has prograded during the past century as the result of sand accretion, forming an outer barrier in front of an elongated lagoon. To the south, Birkdale Sands and Ainsdale Sands are backed by grassy dunes on which famous golf courses have

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Lancashire, Merseyside and Cheshire

been developed (Hansom et al. 1993). The dunes are of calcareous sand winnowed from the wide intertidal sand flats. The wide shore has ridge and runnel topography (Gresswell 1953) with ridges and swales (amplitude up to 1.2 m) parallel to the coastline. There has been ­progradation between Southport and Formby Point, where the coast turns SE, but beyond Formby Point the seaward margin of the coastal dunes has been cliffed by erosion. This erosion has extended northward.

Much of the coastline at Crosby, Bootle and the lower Mersey estuary is artificial, with port, industrial and urban development, and the same is true of the east coast of the Wirral Peninsula. Northward the Mersey cuts through the intertidal sandflats of Liverpool Bay as it opens to the Irish Sea, and the north coast of the Wirral Peninsula is fronted by sandy beaches and wide intertidal sand and mudflats (>Fig. 7.3.3). Red Rocks (Hilbre Point) are a Triassic sandstone outcrop, marking a sharp change to the SW facing ⊡⊡ Fig. 7.3.2 Salt marsh bordering the River Lune at Glasson, looking NW to the power station at Heysham. (Courtesy Geostudies.)

⊡⊡ Fig. 7.3.3 The shore at Hoylake, Wirral Peninsula. (Courtesy Geostudies.)

The Isle of Man

7.4

coastline of the Wirral Peninsula. The shore is sandy, backed by dunes and very extensive intertidal sand and muddy sand areas are exposed at low tide besides the River Dee.

i­ ntertidal zone is predominantly muddy in the south, passing northward to mudflats.

4.  Cheshire

Gresswell RK (1953) Sandy shores of South Lancashire. University of Liverpool, Liverpool, UK Hansom JD, Comber DPM, Fahy FM (1993) Ribble estuary. Coastal Pro­ cesses and Conservation Report, University of Glasgow, Glasgow, UK Marker ME (1967) The Dee estuary: its progressive silting and salt marsh development. Trans Inst Brit Geogr 41:65–71

There are extensive salt marshes on the short Cheshire coast, towards the mouth of the River Dee (Marker 1967) and deposition, and within the wide Dee estuary the

References

7.4  The Isle of Man The Isle of Man trends NE–SW and has an interior upland with valleys radiating from Snaefell (620 m), a transverse vale between Douglas and Peel, and a southern upland rising to South Barrule (477 m). In the south the hill country subsides to a low limestone plateau around Castletown, and in the north is the broad lowland of Ayre, composed of Pleistocene glacial drift deposits. The island is dominated by strongly folded Upper Cambrian Manx (Barrule) Slates, with a Caledonian (NE–SW) trend, and includes small areas of Carboniferous Limestone in the south, some intrusions of basalt and diorite, and Permian sandstones near Peel. Permian rocks occupy subsided basins in the Irish Sea, the Isle of Man being an uplifted horst within the Irish Sea basin. In the Pleistocene the Isle of Man was at times covered by an ice sheet moving south across the Irish Sea basin. The glacial drift in the north of the Isle of Man was part of an extensive area of drift deposits across what is now the floor of the Irish Sea. When the ice finally melted there was an episode of glacial readvance, marked by a push moraine in the Bride Hills. Submerged glacial drift, worked over by wave action during and after the Late Quaternary marine transgression, has yielded a vast quantity of gravel, sand and finer sediment to the shores of the Irish Sea. Late in Pleistocene times the Isle of Man was strongly periglaciated, and the most obvious legacy of this is the slope-over-wall profile that occurs around much of the coast. The slope was formed by periglacial degradation of a Pleistocene cliffed coast, and mantled with solifluction (Head) deposits, while the basal cliff has been cut in Holo­cene times, since the Late Quaternary marine transgression brought the sea to about its present level. Apart from the northern peninsula, which is beachfringed with cliffs cut into glacial drift deposits, much of the Isle of Man coast is rocky, with only limited beaches.

Erosion has been prevalent on much of the coastline, ­especially on the soft glacial drift (Rouse 1990). Features of the coastline are described by Evans (1987). At Douglas in the SE of the island the esplanade is fronted at low tide by a wide beach of sand with some grey sandstone and quartz pebbles above high tide level, widening northward in front of the curved sea wall. To the north a steep slope-over-wall coast extends round to Onchan Head. Several streams flow down to the east coast from the interior uplands, and have incised valleys that open on to the steep coast, such as the deep Groudle Glen opening to Port Groudle. Between Clay Head and Laxey steep grassy coastal slopes on Head deposits descend to shore outcrops of grey Cambrian Manx Slates (>Fig. 7.4.1). The steep coast continues north to Ramsey, where an escarpment of Cambrian slates runs inland across to Ballaugh and Kirkmichael. This was an early Pleistocene coast, cliffed before the glaciers lay down the deposits to the north, and subject to intense periglaciation when the ice edge readvanced to form the Bride push-moraine. North of Ramsey the coast has cliffs cut in glacial drift (>Fig. 7.4.2) extending to the Point of Ayre at the northern end of the Isle of Man. This stands on a heathy plain dotted with gravel pits, which incorporates a raised beach. A broad and high gravel beach of shingle and sand wraps round the Point, and has multiple berms. On the NW coast sand and gravel beaches (>Fig. 7.4.3) backed by dunes with heath vegetation extend south towards Jurby, where they give place to cliffs cut in glacial drift (>Fig. 7.4.4). Near Kirkmichael the cliffs are cut in glacifluvial sands (>Fig. 7.4.5), fronted by a sand and gravel beach with some glacial erratic boulders on the shore. To the south the cliffs of glacial drift give place to cliffs in hard rock formations as the north-facing escarpment reaches the coast.

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The Isle of Man

7.4

coastline of the Wirral Peninsula. The shore is sandy, backed by dunes and very extensive intertidal sand and muddy sand areas are exposed at low tide besides the River Dee.

i­ ntertidal zone is predominantly muddy in the south, passing northward to mudflats.

4.  Cheshire

Gresswell RK (1953) Sandy shores of South Lancashire. University of Liverpool, Liverpool, UK Hansom JD, Comber DPM, Fahy FM (1993) Ribble estuary. Coastal Pro­ cesses and Conservation Report, University of Glasgow, Glasgow, UK Marker ME (1967) The Dee estuary: its progressive silting and salt marsh development. Trans Inst Brit Geogr 41:65–71

There are extensive salt marshes on the short Cheshire coast, towards the mouth of the River Dee (Marker 1967) and deposition, and within the wide Dee estuary the

References

7.4  The Isle of Man The Isle of Man trends NE–SW and has an interior upland with valleys radiating from Snaefell (620 m), a transverse vale between Douglas and Peel, and a southern upland rising to South Barrule (477 m). In the south the hill country subsides to a low limestone plateau around Castletown, and in the north is the broad lowland of Ayre, composed of Pleistocene glacial drift deposits. The island is dominated by strongly folded Upper Cambrian Manx (Barrule) Slates, with a Caledonian (NE–SW) trend, and includes small areas of Carboniferous Limestone in the south, some intrusions of basalt and diorite, and Permian sandstones near Peel. Permian rocks occupy subsided basins in the Irish Sea, the Isle of Man being an uplifted horst within the Irish Sea basin. In the Pleistocene the Isle of Man was at times covered by an ice sheet moving south across the Irish Sea basin. The glacial drift in the north of the Isle of Man was part of an extensive area of drift deposits across what is now the floor of the Irish Sea. When the ice finally melted there was an episode of glacial readvance, marked by a push moraine in the Bride Hills. Submerged glacial drift, worked over by wave action during and after the Late Quaternary marine transgression, has yielded a vast quantity of gravel, sand and finer sediment to the shores of the Irish Sea. Late in Pleistocene times the Isle of Man was strongly periglaciated, and the most obvious legacy of this is the slope-over-wall profile that occurs around much of the coast. The slope was formed by periglacial degradation of a Pleistocene cliffed coast, and mantled with solifluction (Head) deposits, while the basal cliff has been cut in Holo­cene times, since the Late Quaternary marine transgression brought the sea to about its present level. Apart from the northern peninsula, which is beachfringed with cliffs cut into glacial drift deposits, much of the Isle of Man coast is rocky, with only limited beaches.

Erosion has been prevalent on much of the coastline, ­especially on the soft glacial drift (Rouse 1990). Features of the coastline are described by Evans (1987). At Douglas in the SE of the island the esplanade is fronted at low tide by a wide beach of sand with some grey sandstone and quartz pebbles above high tide level, widening northward in front of the curved sea wall. To the north a steep slope-over-wall coast extends round to Onchan Head. Several streams flow down to the east coast from the interior uplands, and have incised valleys that open on to the steep coast, such as the deep Groudle Glen opening to Port Groudle. Between Clay Head and Laxey steep grassy coastal slopes on Head deposits descend to shore outcrops of grey Cambrian Manx Slates (>Fig. 7.4.1). The steep coast continues north to Ramsey, where an escarpment of Cambrian slates runs inland across to Ballaugh and Kirkmichael. This was an early Pleistocene coast, cliffed before the glaciers lay down the deposits to the north, and subject to intense periglaciation when the ice edge readvanced to form the Bride push-moraine. North of Ramsey the coast has cliffs cut in glacial drift (>Fig. 7.4.2) extending to the Point of Ayre at the northern end of the Isle of Man. This stands on a heathy plain dotted with gravel pits, which incorporates a raised beach. A broad and high gravel beach of shingle and sand wraps round the Point, and has multiple berms. On the NW coast sand and gravel beaches (>Fig. 7.4.3) backed by dunes with heath vegetation extend south towards Jurby, where they give place to cliffs cut in glacial drift (>Fig. 7.4.4). Near Kirkmichael the cliffs are cut in glacifluvial sands (>Fig. 7.4.5), fronted by a sand and gravel beach with some glacial erratic boulders on the shore. To the south the cliffs of glacial drift give place to cliffs in hard rock formations as the north-facing escarpment reaches the coast.

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The Isle of Man

⊡⊡ Fig. 7.4.1 Grassy slopes on Head deposits north of Laxey on the east coast of the Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.2 Cliffs cut in glacial drift north of Ramsey. (Courtesy Geostudies.)

The Isle of Man

⊡⊡ Fig. 7.4.3 Cliffed dunes and zones of sand and shingle on the beach at Lhen, Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.4 Cliffs cut in glacial drift at Jurby Head, Isle of Man. (Courtesy Geostudies.)

7.4

363

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7.4

The Isle of Man

⊡⊡ Fig. 7.4.5 Cliffs in fluvioglacial sands near Kirkmichael, Isle of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.6 Fleshwick Bay. (Courtesy Geostudies.)

The west coast north and south of Peel has steep slopes with basal outcrops of Cambrian rock and bold cliffs on headlands. There are some small bays with beaches of sand and gravel (>Fig. 7.4.6). Port Erin is a wide valley-mouth bay between hogsback cliffs, with quartz-veined phyllites on rocky shores and a sandy beach backed by a sea wall. The steep slopeover-wall coast continues round the southern promontory to Spanish Head. The Calf of Man (>Fig. 7.4.7) is an

­ utlying island separated by Calf Sound, where strong o tidal currents swirl on either side of the rocky iset of Kitterland. Numerous basaltic dykes intersect cliff and shore outcrops, some weathered out as clefts. In the SE of the Isle of Man Carboniferous Limes­­ tone  outcrops on cliffed peninsulas, as at Scarlett Point. Castletown Bay is wide and south-facing, with mainly limestone shores. To the north Santon Head is an outcrop of Carboniferous grits, and the Cambrian Manx Slates come to the steep coast. The coastal slopes are interrupted

The Isle of Man

7.4

⊡⊡ Fig. 7.4.7 View across Calf Sound and Kitterland to The Calf of Man. (Courtesy Geostudies.)

⊡⊡ Fig. 7.4.8 Steep coast south of Douglas, Isle of Man. (Courtesy Geostudies.)

at Port Soderick, where a cove stands at the mouth of Glen Soderick, a deeply incised valley. The steep coast on strongly folded Cambrian sandstones and shales continues from Little Ness to Douglas Head (>Fig. 7.4.8), the steep slopes mantled by earthy Head  deposits. There are coves, clefts (zawns) and deep inlets in shales and buttresses on the harder sandstones. Landslides have disrupted the Marine Drive towards Douglas Head, where vertical and steeply dipping slates show breakaways along bedding planes. Landslide debris

has supplied grey shingle to pocket beaches in coves, and there is a natural arch in the cove below the largest slide.

References Evans A (1987) Isle of Man coastal path. Cicerone, Milnthorpe, Cumbria, UK Rouse C (1990) The Isle of Man’s unstable coast. In: Robinson V, McCarroll D (eds) The Isle of Man: celebrating a sense of place. Liverpool University Press, Liverpool, UK

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7.5

North Wales and Anglesey

7.5  North Wales and Anglesey 1.  Introduction The geological formations of North and West Wales show the Caledonian trend (NE–SW), well displayed in the grain of Anglesey and the projection of the Lleyn Peninsula, but complicated by the Harlech (Cambrian) Dome. The geological strike curves westward behind Cardigan Bay and meets the Armorican trend (E–W) in Pembrokeshire and extending through South Wales. The interior includes several planation surfaces, shaped by erosion in relation to higher sea levels in Tertiary and Quaternary times (Brown 1960). There is an early Pleistocene coastline at about 180 m, and a series of marine terraces at lower ­levels, notably between 45 and 75 m, prominent in Pembrokeshire. Glacial drift deposits are extensive, and in coastal regions where they have been cliffed by marine erosion they have yielded sand and gravel to beaches and tidal sand flats and finer sediment to salt marshes and tidal mudflats. Beaches have also received sand and gravel swept in from the sea floor by wave action, but fluvial sediment yields have generally been too fine for retention on beaches. Much of the coast is cliffed, with interruptions at the mouths of valleys that were submerged by the Late Quaternary marine transgression and have been to varying degrees infilled with sediment brought down by rivers and swept in from the sea.

2.  North Wales and Anglesey The north coast of Wales has been much influenced by  the  deposition of Pleistocene glacial drift deposits, which largely conceal the solid rock formations that outcrop in the hinterland. Carboniferous Limestone comes to the coast west of Abergele, and is succeeded by Lower Palaeozoic volcanic rocks west of Conway. Anglesey has parallel outcrops of Pre-Cambrian, Ordovician and Car­ boniferous formations, trending NE–SW, as in Menai Strait which separates it from the mainland. Mean spring tide range on the north coast of Wales is generally between 6 and 7 m (Llandudno 6.9 m, Trwyn Du 6.9 m, Menai Bridge 6.6 m and Amlwch 6.4 m), but diminishes round Anglesey (Wylfa Head 5.7 m, Holyhead 4.9 m, Porth Rhuffydd 4.3 m, Llanddwyn Island 3.9 m). The Point of Air in NE Wales is a foreland of grassy dunes over shingle. To the west a wide alluvial coastal plain is backed by steep hills at the northern end of the Clwydian Range. Wide sand flats are exposed at low tide. Between Prestatyn and Rhyl erosion has long been a

­ roblem, and much of the coastline is armoured by sea p walls and boulder ramparts. Longshore drifting is predominantly eastward along the coast, and a spit that has grown east from Kinmel Bay deflecting the mouth of the River Clwyd. Towards Abergele the sandy shore is backed by a narrow coastal dune fringe. The coastal plain narrows westward in front of a steep slope cut in hard Silurian and Lower Carboniferous rocks, which comes to the coast at Colwyn Bay. The coast then curves out to a promontory of  Carboniferous Limestone at Rhos-onSea. To the west the land rises to Little Ormes Head, also Carboniferous Limestone, with a 141 m summit and much bare rock. Llandudno has a promenade with a high wide beach of grey shingle declining westward to sand as the coast swings out to limestone cliffs of Great Ormes Head. This is a massive, cliff-edged block of Carboniferous Limestone, rising to a summit of 207 m and linked to mainland by an isthmus of dune sand over glacial drift. It is an elongated synclinal structure with an axis SW–NE. The northern coast is an escarpment on dipping limestone (>Fig. 7.5.1). To the south is Conwy Bay, which at low tide has a  wide expanse of sandflats and mudflats with bars and  tidal channels, and sandy beaches on either side of  the  Conwy estuary. West of Conwy sharply rising slopes on Ordovi­cian shales and grits run out to a steep promontory at Penmaen-bach Point. At Penmaenmawr (>Fig.  7.5.2) a steep slope of rhyolite rises behind a coastal fringe of glacial drift, and has been extensively quarried. There has been slope stabilisation above the coast road and railway here, with contour fencing to intercept falling rocks. Much of the coastline is artificial, with defensive walls. The intertidal zone widens rapidly to the extensive Lavan Sands, an area of sand and mud (reworked glacial drift). Seaweed-covered boulders, Llys Helig, extend a mile offshore, uncovered at low tide, the remains of a glacial moraine from which the finer material was swept away by the sea during the later stages of the Late Quaternary marine transgression. Menai Strait is 20 km long, and was formed by the submergence of valleys that diverged from a divide that is now submerged. It is a drowned lowland following the NE–SW geological strike, and is now sluiced by strong tidal currents. On the northern side of the strait a steep wooded coast runs eastward past the foreland of Gallows Point and along

North Wales and Anglesey

7.5

⊡⊡ Fig. 7.5.1 Eastward-dipping Carboniferous Limestone at St Tudwal’s Buttresses on Great Ormes Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.5.2 The steep rhyolite slope west of Penmaenmawr. (Courtesy Geostudies.)

to Beaumaris. Towards Beaumaris the beach becomes sandy, and drifting sand has accreted alongside the breakwater. A succession of bays and minor headlands with cliffs cut into drumlins and fluvioglacial deposits then extends north to Penmon, and off Black Point is Puffin Island, a steep-sided grassy island of Carboniferous Limestone. West from Black Point a steep coast in Carboniferous Limestone extends round to Red Wharf Bay, a 4 km wide bay with extensive sand flats exposed at low tide. It stands at the north-eastern end of the lowland corridor that runs

across the island from Malltraeth Bay. Bluffs and cliffs run round to Benllech. The coast to Moelfre is cliffed and rocky, with a sandy cove at Traeth Bychan. Moelfre stands on a cliffed promontory and cliffs then pass behind long sandy beaches, Traeth Lligwy and Traeth-yr-Ora. The north coast of Anglesey has slope-over-wall cliffs in hard Pre-Cambrian rock (>Fig. 7.5.3), dissected into coves and clefts, headlands and stacks, in relation to variations in resistance of component formations.

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North Wales and Anglesey

⊡⊡ Fig. 7.5.3 Cliffs in Pre-Cambrian gneiss at St. Patrick’s church, Llanbadrig, Anglesey. (Courtesy Geostudies.)

⊡⊡ Fig. 7.5.4 The White Lady, a stack on the east coast of Cemmaes Bay, Anglesey. (Courtesy Geostudies.)

Cemmaes Bay has cliffs cut into Ordovician shales, an  exhumed sub-drift landscape. Two glacial drift-filled ­valleys enter the bay. On the southern coast is White Lady  Beach, a bay with a tall white stack in the middle (>Fig. 7.5.4) and a beach of well-rounded pebbles. At Cemmaes Head the cliffed coast swings south past a series of small coves, declining behind Church Bay, a sandy cove. Sandy beaches alternate with rocky and shore platforms. Holyhead Bay has wide intertidal sand flats on either

side of a channel that leads out from a lagoon. On its western coast a drumlin has been truncated by a marine cliff. Holy Island is separated from Anglesey by an irregular strait running south from Holyhead Bay to a narrow outlet to Cymyran Bay. Inland slopes rise to Holyhead Mountain (220 m). North Stack is a tall cliffy islet of grey strongly-folded quartzite, and rugged cliffs extend behind to South Stack, also of quartzite, linked to the mainland by a suspension bridge.

West Wales

Bold cliffs continue along the south coast of Holy Island to Cymyran Bay, which has a 3 km long sandy beach, wide at low tide, in front of low cliffs and grassy dunes, extending to the rocky reefs at Rhosneigr. Head­ lands and bays alternate to the SE at the ends of ridges and valleys down to Malltraeth Bay, which has 5 km of  sand between the low spring tide line and a high grassy embankment that keeps the sea out of Malltraeth Marsh. Llanddwyn Island, a long rocky peninsula of PreCambrian volcanic rocks, rises to 12 m and is notable for  exposures of pillow lavas in cliff and foreshore outcrops. It  projects southward from a sandy beach and a coast fringed by hummocky dunes. To the east the dunes form Newborough Warren, where three successive ridges of dune sand have migrated inland, with associated

7.6

v­ egetation zones (Ranwell 1958, 1972). On the southern shore is a sand and shingle beach, and eastward drifting has formed a spit with several recurves at Abermenai Point, bordering  the southern entrance to Menai Strait (Robinson 1980).

References Brown EH (1960) The relief and drainage of Wales. University of Wales Press, Cardiff, UK Ranwell DS (1958) Movement of vegetated sand dunes at Newborough Warren, Anglesey. J Ecol 46:83–100 Ranwell DS (1972) Ecology of salt marshes and sand dunes. Chapman & Hall, London Robinson AHW (1980) The sandy coast of south-west Anglesey. Trans. Anglesey Antiq. Field Club for 1980:37–66

7.6  West Wales The NE–SW Caledonian trend is evident in the PreCambrian, Silurian and Ordovician rocks of the Lleyn Peninsula but to the east the Cambrian formations come to the surface in the Harlech Dome, the coastline between Criccieth and Barmouth skirting its western edge. To the south, between Tywyn and Fishguard, the Caledonian trend brings Lower Silurian and Upper Ordovician formations to the coast of Cardigan Bay. West of Fishguard the coast to St David’s Head has Lower Ordovician formations, with numerous igneous intrusions forming headlands between valleys and bays cut in the softer Ordovician rocks. The coasts of Cardigan Bay are dominated by bevelled cliffs, typically with straight slopes descending to a steep or vertical rock wall. There is often a basal terrace of boulder clay, the remains of an extensive boulder clay apron that has been cut back by the sea to form the slope-overwall profiles. Irish Sea ice extended along the coast in late Pleistocene times, impounding proglacial lakes in valleys with the result that overflow channels were cut locally. Mean spring tide ranges are generally about 4 m. Porth Dinnle and Bardsey Island have 3.7 m, Pwllelhi 4.3 m, Porthmadog and Barmouth 4.4 m, Aberdovey and Aberystwyth 4.3 m, New Quay and Port Cardigan 4.1 m and Fishguard 3.9 m. A shingle spit borders the low-lying Morfa Dinnle, south of the entrance to Menai Strait, and Dinas Dinnle is a coastal hill where an Iron Age camp has been truncated by a cliff cut in glacial drift deposits. South from Dinas Dinlle the cliffs are cut in boulder clay at Llandwrog

(>Fig. 7.6.1) and rise past Pontllyfni behind a bouldery shore with some sand and gravel beaches. South of Trefor the coast steepens as hinterland slumping slopes rise to Yr Eifl, and there are coastal quarries in the igneous rocks. Quarry waste has spilled down the very steep slope to the rocky shore (>Fig. 7.6.2). Beyond the rocky headland at Penrhyn Glas an apron terrace of glacial drift descends to cliffs with basal rock outcrops at Porth Pistyll. The Trwyn Porth promontory is of Pre-Cambrian pillow lavas. At Porth Nefin glacial drift occupies a deep hollow excavated by the sea, and a wide sandy beach fronts cliffs and bluffs incised by narrow steep-sided chines (>Fig. 7.6.3). The cliffs and bluffs continue SW past Rhos-y-Llan, punctuated by small coves with sandy beaches and rock pools exposed at low tide. At Porth Colmon (>Fig. 7.6.4) a Pleistocene 6 m pebbly raised beach is well exposed, overlain by glacial drift deposits (Whittow 1960). At Porth Oer cliffs cut in glacial drift interrupt the PreCambrian rocky coast, and there is a long beach with the Whistling Sands, which squeak when walked upon, backing bluffs in glacial drift and a rocky foreshore with pools exposed at low tide. SW of Porth Oer the cliffs are cut into the slopes descending from Mynydd Anelog, and then the lower slopes of Mynydd Mawr. Bardsey Island is of Pre-Cambrian rocks, and intervening Bardsey Sound is a tidal strait scoured to 40 m by strong tidal currents. At the southern end of the Lleyn Peninsula the cliffs turn northward into Aberdaron Bay along a steep and cliffed coast. This bay has been cut out

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Bold cliffs continue along the south coast of Holy Island to Cymyran Bay, which has a 3 km long sandy beach, wide at low tide, in front of low cliffs and grassy dunes, extending to the rocky reefs at Rhosneigr. Head­ lands and bays alternate to the SE at the ends of ridges and valleys down to Malltraeth Bay, which has 5 km of  sand between the low spring tide line and a high grassy embankment that keeps the sea out of Malltraeth Marsh. Llanddwyn Island, a long rocky peninsula of PreCambrian volcanic rocks, rises to 12 m and is notable for  exposures of pillow lavas in cliff and foreshore outcrops. It  projects southward from a sandy beach and a coast fringed by hummocky dunes. To the east the dunes form Newborough Warren, where three successive ridges of dune sand have migrated inland, with associated

7.6

v­ egetation zones (Ranwell 1958, 1972). On the southern shore is a sand and shingle beach, and eastward drifting has formed a spit with several recurves at Abermenai Point, bordering  the southern entrance to Menai Strait (Robinson 1980).

References Brown EH (1960) The relief and drainage of Wales. University of Wales Press, Cardiff, UK Ranwell DS (1958) Movement of vegetated sand dunes at Newborough Warren, Anglesey. J Ecol 46:83–100 Ranwell DS (1972) Ecology of salt marshes and sand dunes. Chapman & Hall, London Robinson AHW (1980) The sandy coast of south-west Anglesey. Trans. Anglesey Antiq. Field Club for 1980:37–66

7.6  West Wales The NE–SW Caledonian trend is evident in the PreCambrian, Silurian and Ordovician rocks of the Lleyn Peninsula but to the east the Cambrian formations come to the surface in the Harlech Dome, the coastline between Criccieth and Barmouth skirting its western edge. To the south, between Tywyn and Fishguard, the Caledonian trend brings Lower Silurian and Upper Ordovician formations to the coast of Cardigan Bay. West of Fishguard the coast to St David’s Head has Lower Ordovician formations, with numerous igneous intrusions forming headlands between valleys and bays cut in the softer Ordovician rocks. The coasts of Cardigan Bay are dominated by bevelled cliffs, typically with straight slopes descending to a steep or vertical rock wall. There is often a basal terrace of boulder clay, the remains of an extensive boulder clay apron that has been cut back by the sea to form the slope-overwall profiles. Irish Sea ice extended along the coast in late Pleistocene times, impounding proglacial lakes in valleys with the result that overflow channels were cut locally. Mean spring tide ranges are generally about 4 m. Porth Dinnle and Bardsey Island have 3.7 m, Pwllelhi 4.3 m, Porthmadog and Barmouth 4.4 m, Aberdovey and Aberystwyth 4.3 m, New Quay and Port Cardigan 4.1 m and Fishguard 3.9 m. A shingle spit borders the low-lying Morfa Dinnle, south of the entrance to Menai Strait, and Dinas Dinnle is a coastal hill where an Iron Age camp has been truncated by a cliff cut in glacial drift deposits. South from Dinas Dinlle the cliffs are cut in boulder clay at Llandwrog

(>Fig. 7.6.1) and rise past Pontllyfni behind a bouldery shore with some sand and gravel beaches. South of Trefor the coast steepens as hinterland slumping slopes rise to Yr Eifl, and there are coastal quarries in the igneous rocks. Quarry waste has spilled down the very steep slope to the rocky shore (>Fig. 7.6.2). Beyond the rocky headland at Penrhyn Glas an apron terrace of glacial drift descends to cliffs with basal rock outcrops at Porth Pistyll. The Trwyn Porth promontory is of Pre-Cambrian pillow lavas. At Porth Nefin glacial drift occupies a deep hollow excavated by the sea, and a wide sandy beach fronts cliffs and bluffs incised by narrow steep-sided chines (>Fig. 7.6.3). The cliffs and bluffs continue SW past Rhos-y-Llan, punctuated by small coves with sandy beaches and rock pools exposed at low tide. At Porth Colmon (>Fig. 7.6.4) a Pleistocene 6 m pebbly raised beach is well exposed, overlain by glacial drift deposits (Whittow 1960). At Porth Oer cliffs cut in glacial drift interrupt the PreCambrian rocky coast, and there is a long beach with the Whistling Sands, which squeak when walked upon, backing bluffs in glacial drift and a rocky foreshore with pools exposed at low tide. SW of Porth Oer the cliffs are cut into the slopes descending from Mynydd Anelog, and then the lower slopes of Mynydd Mawr. Bardsey Island is of Pre-Cambrian rocks, and intervening Bardsey Sound is a tidal strait scoured to 40 m by strong tidal currents. At the southern end of the Lleyn Peninsula the cliffs turn northward into Aberdaron Bay along a steep and cliffed coast. This bay has been cut out

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⊡⊡ Fig. 7.6.1 Boulder clay cliffs at Llandwrog, looking south towards the Yr Eifl mountains. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.2 Quarry waste spilling down coastal slope, Yr Eifl. (Courtesy Geostudies.)

between cliffy promontories of Pre-Cambrian rock to the west and Ordovician to the east, but at the head of the bay the south-facing cliffs are cut in soft boulder clay. On the eastern side of Aberdaron Bay the cliffs run out to a headland at Trwyn-y-Penrhyn, and there is a 6 m raised beach around offshore islands. A cliffed promontory runs past Porth Ysgo, where the coastal slope has been disturbed by former mining, and round beside Porth Neigwl, a wide bay with receding cliffs cut in soft boulder clay along a gently curved coastline, shaped by

refracted south-westerly waves. East from Abersoch is a  succession of asymmetrical embayments with long beaches of sand and gravel between rocky headlands. Extensive intertidal sands at the head of Tremadog Bay, passing into the Glaslyn estuary, are the outcome of northward drifting of beach material from Cardigan Bay and eastward drifting from Lleyn, together with shoreward drifting from the sea floor. The long sandy estuaries of the Glaslyn and the Dwyryd have been much modified by reclamation.

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⊡⊡ Fig. 7.6.3 Porth Nefin. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.4 Emerged beach at Porth Colmon. (Courtesy Geostudies.)

Morfa Harlech is a broad triangular foreland south of the Glaslyn estuary, consisting of an undulating lowland of glacial drift and alluvium, with hollows occupied by salt marshes and freshwater swamps. The western shore is backed by shingle and sand beaches backed by dunes. The cuspate foreland at Morfa Dyffryn is similar to Morfa Harlech in that it consists of glacial drift with a fringe of beaches and dunes, particularly on the SW flank. Both forelands have grown northward as the result of longshore drifting and accretion, but their SW coasts are eroding.

South of Llanaber cliffs and shingle beaches extend to Barmouth, which stands on a small sandy foreland at the mouth of the Mawddach estuary. This estuary occupies a valley cut into Cambrian beds on the southern side of the Harlech Dome, and has been formed by Late Quaternary marine submergence. Ro Wen is a shingle and sand spit that runs north from Fairbourne and almost blocks the mouth of the estuary. South of Fairbourne a steep coast with basal cliffs and a boulder beach is interrupted by a delta-like lobe of glacial drift at

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the mouth of Afon Gwrill. The coastal plain extends southwards past Tywyn. An alluvial coastal plain continues southward, bordered by dunes and a beach of shingle and sand which curves eastward to Aberdovey. The Dyfi estuary is cut in strongly-folded slates, grits and sandstones which strike NNE–SSW. It is bordered by salt marshes with grazed turf, cliffed edges seaward and undercut creek banks (Burd 1989). On the southern side the Borth sand and shingle spit has grown about 3 km northward from Upper Borth, supplied with sand and shingle drifting from the south and shaped by westerly waves. At Upper Borth (>Fig. 7.6.5) a storm beach of grey platy cobbles is backed by a sea wall, and fronted by a wide sandy surf beach with remnants of a submerged forest exposed at low tide, with peat deposits are over estuarine clay. The peat formed about 6,000 years ago, and is exposed because of eastward migration of the spit. The beach ends as cliffs rise behind a shore platform. From Upper Borth to Aberystwyth and New Quay the Silurian Aberystwyth Grits form cliffs and shore platforms, locally mantled by boulder clay and interrupted by valleys floored with glacial drift. Where the dip is landward the cliffs slope seaward, where it is horizontal they are more or less vertical, and where it is seaward there is landslide topography. At the mouth of the Wallog valley there is a mantle of boulder clay, and a gravelly cause­way, Sarn Cynfelyn, well exposed at low tide, runs out into Cardigan Bay.

The cliffs come to an end at Aberystwyth, where a promenade runs behind a beach of shingle and a sandy foreshore exposed at low tide. To the south the Afon Rheidol comes in from the east to meet the Afon Ystwyth on the seaward side of high Pendinas Hill. The Ystw­­ yth flows in from the south across an alluvial plain behind a shingle barrier beach, then the coast rises to cliffs at Peny-bwlch, truncating a coastal ridge. The shore is littered with boulders, apart from a sector of abrasion platform cut across the steeply dipping mudstones (>Fig. 7.6.6). The cliffs to the south are steep and up to 30 m high. They are interrupted by valleys, and in places fronted by a boulder clay apron. A deep, narrow valley opens to the coast at Aberarth, where a 4 m cliff stands behind a bouldery cobble beach and the river mouth has been diverted northward by a small spit. Between Aberarth and Aberaeron there are cliffs cut in boulder clay. At Aberaeron beach shingle drifts northward, and by the early nineteenth century it had deflected the river mouth of the Afon Aeron about 650 m northward, but in 1807 it was diverted back to the harbour. To the south the shingle beach narrows in front of rising cliffs. New Quay stands in a bay with slumping cliffs in soft grey shale, and at New Quay Head the cliffed coastline turns WSW, rising along the flank of a coastal ridge. Another deep narrow valley descends to Cwm Tydi, a cove with a sea wall behind a beach of sand and platy blue– grey cobble gravel, and to the SW is an intricate coastline with coves and small headlands, including the narrow

⊡⊡ Fig. 7.6.5 The upper beach of shingle behind a sandy foreshore at Upper Borth. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.6.6 Shore platform on seaward dipping mudstones at Pen-yBwlch. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.7 Coastal slope west of Cwm Tydi. (Courtesy Geostudies.)

irregular Ynys Lochtyn peninsula (>Fig. 7.6.7) with a 30 m high coastal slope descending to cliffs and clefts. The steep coast continues westward behind Traeth Penbryn, a long beach of yellow sand with some gra­­ vel,  backed by low dunes and cliffs cut in boulderclay. Aberporth is in a double cove where two valleys ­converge, then the coast rises to the high headland above Cribach Bay. The coastal plateau is trenched by ­deeply-incised V-shaped valleys that may have been meltwater channels cut by overflow from an impounded glacial lake.

To the south Traeth-y-Mwnt (>Fig. 7.6.8) is a cove bordered by cliffs cut in steeply-dipping dark grey sedimentary rocks and backed by a 20 m high cliff cut in ­glacial drift and capped by grassy dunes. There is a beach of ­yellow–brown sand, wide at low tide. South from Mwnt the cliffs are rugged, with coves, caves and natural arches. They have slope-over-wall profiles, which also extend round Cardigan Island offshore. A funnel-shaped estuary at the mouth of the River Teifi opens between a broad sandy lobate foreland on the

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⊡⊡ Fig. 7.6.8 Traeth-y-Mwnt. (Courtesy Geostudies.)

⊡⊡ Fig. 7.6.9 Dinas Head. (Courtesy Geostudies.)

s­ outhern shore, the beach backed by grassy dunes, and a dune-capped spit at Gwbert on the northern shore. At low tide the waves break around the outer edge of Poppit Sands, a wide intertidal sand flat with bands of paler yellow sand marking multiple strandlines. Salt marshes have been dissected by tidal creeks and cliffed at the outer edge by wave erosion. Cemmaes Head is backed by overflow channel from a glacial lake that occupied the Teifi valley when ice blocked its mouth. Slope-over-wall cliffs continue south–west to Newport Bay, where Newport Sands are extensive at low

tide, the Afon Nyfer winding across them to the sea between sand spits. Inland the ground rises towards Presceli Mountain. The coast swings out to the high promontory of Dinas Head, rising 142 m above sea level (>Fig. 7.6.9), bordered on its landward side by a valley cut by meltwater overflow from a glacial lake. West from Dinas Head cliffs in dark Ordovician shales and sandstones are broken by deep and narrow valleys. Towards Fishguard the cliffs are on paler igneous rocks, on an indented coast, and the large breakwater at

South Wales

Fishguard Harbour lies under a cliff on the western side of the bay. Strumble Head is a bold heathy promontory with cliffs exposing pillow lavas, rhyolites, agglomerates and volcanic ash. To the south are coves cut in Ordovician shales between steep bluffs, cliffs and rocky shores on volcanic and intruded rocks. There is an extensive capping of boulder clay, and flint gravel is found on beaches. Abercastle is behind a long narrow drowned valley mouth. The sequence of headlands in dolerite or rhyolite and bays and inlets in softer Ordovician sedimentary rocks continues SW to St. David’s Head. As most of the

7.7

rivers on the St. Davids Peninsula rise close to the north coast and flow south there are few sandy coves on the cliffy northern shore.

References Burd F (1989) The saltmarsh survey of Great Britain. Research and Survey in Nature Conservation No. 17, Peterborough Whittow JB (1960) Some comments on the raised beach platform of south-west Carnarvonshire and on an unrecorded raised beach at Porth Neigwl. Proc Geologis Assoc 71:31–39

7.7  South Wales Along the south coast of Wales the geological formations trend generally west to east, and the coastline intersects outcrops of rock formations that are often steeply dipp­­ ing.  The Pre-Cambrian and Lower Palaeozoic rocks of St. David’s Peninsula are succeeded southward in St. Bride’s Bay by Devonian (Old Red Sandstone) and Carboniferous formations, including Carboniferous Limestone, Millstone Grit and Coal Measures. These extend W–E across south Pembrokeshire, Carmarthen Bay and the Gower Peninsula to Swansea Bay and Porthcawl. There are broad plateaux ending in even-crested cliffs, except where incised valleys open to the coast. In Glamorgan the Carboniferous rocks subside beneath Trias and Lias formations. Keuper Marl forms a coastal fringe north–east from Cardiff, but is largely concealed by Pleistocene glacial drift and Holocene alluvium in the Wentlooge and Caldicot Levels. Mean spring tide ranges increase W–E along the coast. On Skomer Island the range is 5.9 m and at Milford Haven 6.3 m. Stackpole Quay has 7.0 m and Tenby 7.6 m. On Gower the range at Whitford Lighthouse is 7.8 m and in Swansea Bay it increases from 8.6 m at Swansea to 8.7 m at Port Talbot and 8.9 m at Porthcaw l8.9. The increase ­continues as the Bristol Channel narrows into the Severn estuary: Barry has 10.7 m, Cardiff 11.1 m, Newport 11.9 m, Sudbrook 12.4 m and Beachley, at the Severn Bridge, 12.4 m. In Pleistocene times the Devensian Glaciation is thought to have extended to South Wales, the limit (c 20,000 bp) of Irish Sea ice extending from the southern side of St. Brides Bay inland over north Pembrokeshire to meet Welsh Ice that extended back out across Carmarthen Bay and the northern part of the Gower Peninsula, across Swansea Bay to Porthcawl, then inland behind South Glamorgan to a lobe between Cardiff and Newport.

On the Pembrokeshire coast (John 1979) St. David’s Head is a promontory of gabbro. Cliffs cut in Cambrian rocks fringe the coast into Whitesands Bay, which has a typical west-facing Atlantic surf beach, wide at low tide and backed by dunes. Clefts and caves have been etched out along faults in the Cambrian rocks along the shore, and there is an emerged shore platform with a raised beach of sand and shingle. The St. David’s Peninsula is ­dominated by a broad plateau 45–75 m above sea level, generally ­considered to be a marine planation surface with residual higher hills that were once islands when the sea stood at a higher level. The Bitches is a tide race inshore from Ramsey Island, with strong currents as the tide rises and falls. Ramsey Island is a detached slab of the coastal plateau bordered by  cliffs in Ordovician shales between headlands of quartz-porphyry. The mainland coast swings eastward with slope-over-wall cliffs in Pre-Cambrian volcanic rocks, with intruded granites and dolerites forming headlands and islands. East of Pen-y-Cyrfrwy a coastal fringe of Cambrian rocks outcrops in the cliffs, and these continue past Caerfai Bay to the Solva inlet, and on to the mouth of the Cwm-mawr valley, where a major fault marks the beginning of the strongly folded Carboniferous formations (Coal Measures) which dominate the east coast of St. Bride’s Bay. At Newgale the sandy beach, wide at low tide, is backed by a straight ridge of storm-piled grey and brown cobbles. The cliffs are cut in Coal Measure Sandstones and Millstone Grit, overlain by glacial drift. Haroldston Bridge is a natural arch under a dipping ramp of Coal Measures sandstone, and blocks of cliff are falling away at Haraldston Chins, leaving deep chasms. Below the cliffs to the south is the Sleek Stone (>Fig. 7.7.1), a monoclinal protrusion of

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Fishguard Harbour lies under a cliff on the western side of the bay. Strumble Head is a bold heathy promontory with cliffs exposing pillow lavas, rhyolites, agglomerates and volcanic ash. To the south are coves cut in Ordovician shales between steep bluffs, cliffs and rocky shores on volcanic and intruded rocks. There is an extensive capping of boulder clay, and flint gravel is found on beaches. Abercastle is behind a long narrow drowned valley mouth. The sequence of headlands in dolerite or rhyolite and bays and inlets in softer Ordovician sedimentary rocks continues SW to St. David’s Head. As most of the

7.7

rivers on the St. Davids Peninsula rise close to the north coast and flow south there are few sandy coves on the cliffy northern shore.

References Burd F (1989) The saltmarsh survey of Great Britain. Research and Survey in Nature Conservation No. 17, Peterborough Whittow JB (1960) Some comments on the raised beach platform of south-west Carnarvonshire and on an unrecorded raised beach at Porth Neigwl. Proc Geologis Assoc 71:31–39

7.7  South Wales Along the south coast of Wales the geological formations trend generally west to east, and the coastline intersects outcrops of rock formations that are often steeply dipp­­ ing.  The Pre-Cambrian and Lower Palaeozoic rocks of St. David’s Peninsula are succeeded southward in St. Bride’s Bay by Devonian (Old Red Sandstone) and Carboniferous formations, including Carboniferous Limestone, Millstone Grit and Coal Measures. These extend W–E across south Pembrokeshire, Carmarthen Bay and the Gower Peninsula to Swansea Bay and Porthcawl. There are broad plateaux ending in even-crested cliffs, except where incised valleys open to the coast. In Glamorgan the Carboniferous rocks subside beneath Trias and Lias formations. Keuper Marl forms a coastal fringe north–east from Cardiff, but is largely concealed by Pleistocene glacial drift and Holocene alluvium in the Wentlooge and Caldicot Levels. Mean spring tide ranges increase W–E along the coast. On Skomer Island the range is 5.9 m and at Milford Haven 6.3 m. Stackpole Quay has 7.0 m and Tenby 7.6 m. On Gower the range at Whitford Lighthouse is 7.8 m and in Swansea Bay it increases from 8.6 m at Swansea to 8.7 m at Port Talbot and 8.9 m at Porthcaw l8.9. The increase ­continues as the Bristol Channel narrows into the Severn estuary: Barry has 10.7 m, Cardiff 11.1 m, Newport 11.9 m, Sudbrook 12.4 m and Beachley, at the Severn Bridge, 12.4 m. In Pleistocene times the Devensian Glaciation is thought to have extended to South Wales, the limit (c 20,000 bp) of Irish Sea ice extending from the southern side of St. Brides Bay inland over north Pembrokeshire to meet Welsh Ice that extended back out across Carmarthen Bay and the northern part of the Gower Peninsula, across Swansea Bay to Porthcawl, then inland behind South Glamorgan to a lobe between Cardiff and Newport.

On the Pembrokeshire coast (John 1979) St. David’s Head is a promontory of gabbro. Cliffs cut in Cambrian rocks fringe the coast into Whitesands Bay, which has a typical west-facing Atlantic surf beach, wide at low tide and backed by dunes. Clefts and caves have been etched out along faults in the Cambrian rocks along the shore, and there is an emerged shore platform with a raised beach of sand and shingle. The St. David’s Peninsula is ­dominated by a broad plateau 45–75 m above sea level, generally ­considered to be a marine planation surface with residual higher hills that were once islands when the sea stood at a higher level. The Bitches is a tide race inshore from Ramsey Island, with strong currents as the tide rises and falls. Ramsey Island is a detached slab of the coastal plateau bordered by  cliffs in Ordovician shales between headlands of quartz-porphyry. The mainland coast swings eastward with slope-over-wall cliffs in Pre-Cambrian volcanic rocks, with intruded granites and dolerites forming headlands and islands. East of Pen-y-Cyrfrwy a coastal fringe of Cambrian rocks outcrops in the cliffs, and these continue past Caerfai Bay to the Solva inlet, and on to the mouth of the Cwm-mawr valley, where a major fault marks the beginning of the strongly folded Carboniferous formations (Coal Measures) which dominate the east coast of St. Bride’s Bay. At Newgale the sandy beach, wide at low tide, is backed by a straight ridge of storm-piled grey and brown cobbles. The cliffs are cut in Coal Measure Sandstones and Millstone Grit, overlain by glacial drift. Haroldston Bridge is a natural arch under a dipping ramp of Coal Measures sandstone, and blocks of cliff are falling away at Haraldston Chins, leaving deep chasms. Below the cliffs to the south is the Sleek Stone (>Fig. 7.7.1), a monoclinal protrusion of

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⊡⊡ Fig. 7.7.1 The Sleek Stone, St. Bride’s Bay. (Courtesy Geostudies.)

Coal Measures sandstone, steeper and eroded on its northern flank, the upper surface of a sandstone layer from which overlying softer rocks have been removed. Broad Haven has been cut into strongly-folded Coal Measures sandstones, with smooth bedding planes exposed in the cliffs and Little Haven has been cut into soft shales between sandstone headlands. At Falling Cliff, where subsidence is due partly to the collapse of old coal mineshafts, the Coal Measure sandstones come to an end against dolerite. The coastline swings westward and the cliffs pass into scrubby bluffs on Coal Measures with bolder headlands on Pre-Cambrian volcanic rocks. There are several steepsided coves, as at Foxes Holes, where the Pre-Cambrian rocks give place to Old Red Sandstone, with inlets along small faults as at Musselwick Sands. Beyond Martins Haven is a broad promontory of igneous rocks, Deer Park, with Wooltack Point a headland at the NW corner, and The Anvil to the SE. The western cliffs look across a strait where the tidal currents are strong to Midland Isle, then another strait to Skomer Island. This island is mainly composed of volcanic rocks exposed in cliffs, and Grassholm Island, out to the west, is a cliffy island of basalt. Cliffs on the south coast of Deer Park are strongly influenced by the structure of outcrops of Silurian slate (>Fig. 7.7.2), and the coast runs out to a small headland behind Gateholm Island. This flat-topped island on steeply-dipping Old Red Sandstone is attached to the coast by a rocky reef. Offshore to the SW is Skokholm, a flattopped cliff-edged island of Old Red Sandstone rising 33 m above sea level.

The cliffs run on past Raggle Rocks, and The Three Chimneys are vertically projecting sandstone beds separated by erosion of intervening soft mudstones. Along Marloes Bay the cliffs are cut in colourful strongly folded Silurian rock formations, with bedding planes exposed in the cliff face where the dip is seaward Marloes Beach (>Fig. 7.7.3) is wide and sandy at low tide, when jagged reefs and stacks of Silurian rock protrude from the sand. ­Slope-over wall profiles extend south to St. Ann’s Head and round to Dale Fort. To the east the coast runs along the northern side of Milford Haven, with cliffs in Old Red Sandstone and Silurian rocks. Sandy Haven is a wide inlet at low tide with green weed and brown wrack on gravel and mud. Red Marls underlie the south-facing slopes west and east of the town of Milford Haven. Milford Haven is a large branching ria formed by Holocene marine submergence of the Cleddau River and its tributaries, incised into the Pembrokeshire plateau. In the Cresswell River broad sloping tidal mudflats fringe the low tide channel, and there are grassy salt marshes (>Fig. 7.7.4). On the south coast of Milford Sound coastal slopes descend to low cliffs and rocky shores. Angle Bay is sheltered from strong wave action, and has low tide rocky outcrops and gravel with much seaweed, descending to mudflats. The shore is fringed by salt marsh, dissected by tidal creeks. At the western end of the Angle valley is West Angle Bay, a synclinal cove in Carboniferous Limestone. Slopeover-wall cliffs run out to a headland at East Blockhouse, cut first in Carboniferous Limestone then Old Red

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⊡⊡ Fig. 7.7.2 Cliff in dipping Silurian slates at Horse Neck, near Gateholm, with Skomer Island in the distance. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.3 Marloes Sands. (Courtesy Geostudies.)

Sandstone rising on the southern side of the Angle ­syncline. The cliffs continue to Freshwater West, where Silurian rocks emerge from beneath the Old Red Sandstone as the shore platform disappears beneath the wide Broomhill Sands. Dunes and bluffs facing WSW back a wide sandy shore receiving Atlantic surf and dunes overlie a cobble beach. To the south Old Red Sandstone dips beneath Carboniferous rocks on Limney Head, where the coastline turns ESE.

A very fine stretch of Carboniferous Limestone coast extends ESE from Linney Head, and is well known for its geological structures and the cliff scenery formed by dissection along joints and bedding planes (John 1973). It includes a prominent natural arch at the Green Bridge of Wales (>Fig. 7.7.5) and the outlying Elegug stacks (>Fig. 7.7.6). The cliffs continue eastward to St. Govan’s Head (>Fig. 7.7.7), a bold promontory that shows high vertical

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⊡⊡ Fig. 7.7.4 Salt marsh and tidal mudflats in Cresswell River, Milford Haven. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.5 The Green Bridge of Wales. (Courtesy Geostudies.)

cliffs dissected along rectangular joints, so that blocks of ­limestone have fallen to the shore. Long Matthew Point marks the end of the south-facing cliffed coast. The coastline turns northward and New Quay is a sanded inlet at low; at high tide it is much like a Mediterranean calanque. Behind Broad Haven two valleys incised into the limestone plateau converge east of Bosherston. They were formerly tidal inlets, but in the eighteenth century they were enclosed by a dam as

Bosherston Lakes. The dunes of Stackpole Warren have drifted across from Broad Haven. The cliffs from Stackpole Point north to Trewent Point show the effects of struc­­ ture  and stratigraphy, with sloping ramps on dipping ­bedding-planes, hollows and valleys on shales, and clefts cut out along joints and bedding planes. To the north is Freshwater East Bay, with a wide sandy beach backed by dunes emplaced by easterly winds. It occupies an anticlinal flexure, with Silurian rocks

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⊡⊡ Fig. 7.7.6 Elegug Rocks. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.7 Cliffs of Carboniferous Limestone extending east to St. Govan’s Head. (Courtesy Geostudies.)

­ ipping steeply away on either side of a central outcrop of d Ordovician rocks. The slope-over-wall coast to the east is on steeply-dipping Old Red Sandstone, which in places stand almost vertical, and strike parallel to the coast. Swanlake Bay is a fault-bounded re-entrant with a broad shore platform cut across steeply-dipping Old Red Sandstone strata, also seen at Manorbier (>Fig. 7.7.8). In Skrinkle Haven are vertical strata from the upper Old Red Sandstone to Carboniferous Limestone, which outcrops along the steep coast eastward to Giltar Point

and on outlying Calder Island. Limestone cliffs run northward past Tenby, with Millstone Grit and Coal Measure sandstones towards Monkstone Point. Towards Saundersfoot the coast consists of wooded bluffs with basal cliffs of varying height in folded and faulted Coal Measures, and the beach includes ballast dumped by ships that came in to collect coal. East of Amroth the underlying Millstone Grit outcrops in cliffs that run out to Telpyn Point, and the Carboniferous Limestone forms cliffs along to Dolwen Point. At Pendine

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South Wales

⊡⊡ Fig. 7.7.8 Shore platform cut across steeply dipping Old Red Sandstone in Manorbier Bay. (Courtesy Geostudies.)

these cliffs have several high caves have been cut out along vertical joints in the limestone (>Fig. 7.7.9). Pendine Sands, up to 800 m wide at low tide, border a dune-capped foreland that is backed by bluffs that extend to Ginst Point (Savigear 1952). This shelters a salt marsh beside the sandy Taf estuary, which meets the Tywi and Gwendraeth estuaries in an area of inwashed intertidal sand in the NE of Carmarthen Bay. South of the Gwendraeth estuary these sands are backed by a wide sandy foreland that includes dunes at Tywyn Burrows and Pembrey Burrows. To the east the coast is low and marshy to Burry Port, beside the wide Loughor estuary north of the Gower Peninsula. On its southern coast is the wide Llanrhidian salt marsh (>Fig. 7.7.10) with winding creeks and scattered pools and salt pans, some of which formed as the result of sub-surface tunnelling. The Gower Peninsula consists of folded and faulted Carboniferous Limestone with synclinal furrows occupied by Millstone Grit shales forming bays and anticlinal ridges forming headlands where Old Red Sandstone outcrops. On the west coast, south of the spit at Whitford Burrows, Burry Holms is an island of limestone seaward of a sandy intertidal isthmus. Rhossili Bay has a beach 5 km long, the northern part backed by dunes, the southern part by a gravel beach, low receding cliffs cut into an apron of glacial drift and a gently sloping terrace that rises to steep slopes of Old Red Sandstone on Rhossili Down (>Fig. 7.7.11). Worms Head is a long narrow promontory with a rugged island of strongly folded Carboniferous Limestone at the western end of an intertidal boulder-strewn shore platform.

The SW coast of the Gower Peninsula has steep escarpment cliffs in landward-dipping Carboniferous Limestone rising to a broad plateau. There are there are many minor bays and indentations related to faults or major joints, and segments of 4–5 m raised beach. Paviland Cave is one of several Palaeolithic occupation sites with animal bones and prehistoric implements. A high limestone promontory runs out to flattopped Port Eynon Point and Port Eynon occupies a broad synclinal bay cut in shales along a syncline of Millstone Grit. Oxwich Point is a 75 m high promontory with cliffs on the southern side passing into wooded bluffs beside Oxwich Bay to the north. This has been cut in soft Millstone Grit shales, and has a 2 km wide gentle intertidal sandy shore. On the northern side Great Tor, a barnacled rock pinnacle, protrudes at the western end of Three Cliffs Bay (>Fig. 7.7.12), which is bordered by limestone cliffs and bluffs and has a wide beach backed by grassy dunes. Pwll-du Head is a high (225 m) limestone promontory. To the east Pwll-ddu Bay has been cut out along fault lines, and contains shingle beach ridges. A succession of cliffy headlands and small bays extends to Mumbles Head. Swansea Bay has been cut in soft Upper Carboniferous rocks at the submerged mouths of several ­confluent river valleys. East of the Tawe estuary Kilvey Hill (171 m) rises steeply behind a narrow coastal plain, and sandy beaches backed by dunes extend to Porthcawl. At Porthcawl beach depletion has been countered by making a tarmac shore with embedded cobbles.

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7.7

⊡⊡ Fig. 7.7.9 Caves in Carboniferous Limestone in the cliff at Pendine. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.10 Llanrhidian Marsh. (Courtesy Geostudies.)

Newton Point is a headland projecting southward, and beyond it a curving bay with a broad sandy shore is backed by the grassy dunes of Newton Burrows and Merthyr Mawr. The beach ends at the mouth of Ogmore River, and at Ogmore-by-sea the coast consists of a steep grassy escarpment in thinly-bedded Lias with an apron of  periglacial Head, descending to cliffs cut in the layered  limestone. To the east Carboniferous Limestone strata form sloping stepped ledges, and at Southern­­ down an almost vertical cliff in gently dipping Lias rests

c­ onformably upon Carboniferous Limestone which runs across the shore. At  Nash Point (>Fig. 7.7.13) the cliffs are cut in layered Lias limestone and shales are fronted at low tide by shore platforms up to 500 m wide, with swirling, undulating limestone ledges separated by sinuous scarplets, dissected along joints and bedding planes l (Trenhaile 1972). The limestone cliffs continue SE past Summerhouse Point to Aberthaw and Barry Island (Williams et al. 1993). l Sully Island is a grassy and scrubby island of Triassic rocks

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South Wales

⊡⊡ Fig. 7.7.11 Rhossili Bay, Gower Peninsula. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.12 Three Cliffs Bay, Gower Peninsula. (Courtesy Geostudies.)

with bordering low cliffs and limestone ledges with beach gravels. The low Lias limestone cliffs continue to Lavernock Point, where the coast swings sharply northward to Penarth, crossing an outcrop of Keuper Marl in the core of the Penarth anticline. Off Lavernock Point are Flat Holm and Steep Holm, islands of Carboniferous Limestone in the wide mouth of the Severn estuary.

To the east is the wide mouth of the Taff estuary opening into Cardiff Bay, and the extensive Cardiff docks and waterfront. Beyond the Rhymney estuary is a wide coastal lowland, Peterstone Wentlooge, a former marsh that was drained and reclaimed behind a sea wall. The broad meandering estuary of the River Usk interrupts the coastal plain at Newport, but it resumes to the

South Wales

7.7

⊡⊡ Fig. 7.7.13 Shore platform cut across dipping Lower Lias limestone near Nash Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.14 Cliff in Trias at Black Rock, with fronting salt marsh. (Courtesy Geostudies.)

east as the Caldicot Levels, another area of reclaimed marsh. Much of the coastline consists of a high sea wall, grassy on the inner slope and armoured with boulders on the seaward side, where wide mudflats are exposed at low tide. There are minor cliffs at Sudbrook and Black Rock (>Fig. 7.7.14). At Chepstow the River Wye meanders through a gorge between cliffs cut in massive Carboniferous Limestone before crossing low ground and marshland to open into the Severn estuary west of Beachley Point.

References John BS (1979) The geology of Pembrokeshire. Abercastle Savigear RAG (1952) Some observations on slope development in South Wales. Trans Inst Brit Geogr 31:23–42 Trenhaile AS (1972) The shore platforms of the Vale of Glamorgan. Trans Inst Brit Geogr 56:127–144 Williams AT, Davies P, Bomboe P (1993) Geometrical simulation studies of coastal cliff failures in Liassic strata, South Wales. Earth Surf Process Landforms 18:703–720

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The Severn Estuary

7.7.1  The Severn Estuary The River Severn flows through the wide Vale of Gloucester and downstream from Gloucester widens into the Severn estuary, opening into the Bristol Channel. The South Wales coast east from Penarth shows some estuarine features, as does the Avon coast northeast from Clevedon, but here the shores of the Severn estuary are those upstream from the first Severn Bridge, which crosses from Beachley to Aust. The river is tidal to well above Gloucester, but above Minsterworth the margins are essentially river banks. The Severn estuary is megatidal (mean spring tide range >6 m) and funnel-shaped, narrowing upstream. Mean spring tide range decreases from 12.4 m at Beach­­ ley to 8.8 m at Sharpness Dock and about 2 m at Gloucester. Ebb and flow currents can attain 20 kph, and at spring tides the Severn Bore travels rapidly upstream (>Fig. 7.7.1.1). Wave action is limited by tidal rise and fall, but can be strong with onshore winds on bordering shores at high tide. The present morphology of the Severn estuary results from sediment accumulation in deeper channels that were  cut during Pleistocene phases of low sea level and ­submerged during the Late Quaternary (Flandrian) marine transgression. At low tide, the river banks are fringed by a muddy intertidal slope that widens as the tide range increases downstream. The muddy deposits have been derived from the broad clay vales drained by the Severn and its tributaries, but there has been much

­ ovement of muddy sediment from the Bristol Channel, m brought back upstream by rising tides. Mudflats are shaped by tidal ebb and flow, as well as by wave action as the tides rise and fall (Pethick 1996). The stratigraphy of bordering salt marsh terraces shows alternations of almost horizontal layers of mud interdigitating with backshore peat horizons some of which outcrop as ledges on the seaward slope (Allen 1992).

1.  The North Shore Above Beachley the estuarine shore is fringed by salt  marshes, backed at Sedbury by minor cliffs and steep  wooded bluffs, cut in Rhaetic and Lower Lias (>Fig. 7.7.1.2). The bordering alluvium is interrupted by segments of low cliff at Lydney and Purton, and cliffs cut in thinly bedded grey Triassic limestone and pink shales over red-brown marls towards Awre. Further upstream the meandering river intersects Triassic rocks on the north bank near Newnham.

2.  The South Shore The Severn widens rapidly below Longney, and at Epney the bank at high tide is protected by a stone wall. The river meanders ESE past Hock Cliff, where layers of ⊡⊡ Fig. 7.7.1.1 The Severn Bore passing Stonebench. (Courtesy Geostudies.)

The Severn Estuary

limestone and shale outcrop in the cliff and on the shore (>Fig. 7.7.1.3). At Slimbridge managed retreat is being implemented on land that was formerly embanked and reclaimed. A new clay bank has been built 40 m back from the pre­ vious one, leaving a new intertidal strip to be managed

⊡⊡ Fig. 7.7.1.2 Salt marsh in front of Sedbury Cliff on the Severn shore. (Courtesy Geostudies.)

⊡⊡ Fig. 7.7.1.3 Cracking shales over a ledge of limestone at Hock Cliff on the Severn estuary. (Courtesy Geostudies.)

7. 7.    1

as an estuarine wetland. At Purton the grassy salt marsh ends seaward in a low cliff (>Fig. 7.7.1.4), below which are sloping mudflats with mounds of mud bearing Spartina. Similar features extend downstream, but the shore has  been modified by a large tidal reservoir near Shepperdine.

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Gloucestershire, Somerset and North Devon

⊡⊡ Fig. 7.7.1.4 Erosion of the Severn-side salt marsh terrace at Tites Point, near Purton. (Courtesy Geostudies.)

References Allen JRL (1992) Tidally induced salt marshes in the Severn estuary, southwest Britain. In: Allen JRL, Pye K (eds) Saltmarshes:

morphodynamics, conservation and engineering significance. Cambridge University Press, Cambridge, New York, pp 123–147 Pethick JS (1996) The geomorphology of mudflats. In: Nordstrom KE, Roman CT (eds) Estuarine shores. Wiley, New York, pp 185–211

7.8  Gloucestershire, Somerset and North Devon The coast of Gloucestershire, Somerset and North Devon begins where the Severn estuary runs beside the Vale of Belvoir, a lowland of Keuper Marl. The Jurassic formations of the Cotswolds do not extend to the coast, but there are minor outcrops of Trias at Aust. The Carboniferous Limestone of the Bristol district reaches the seaboard between Portishead and Clevedon (with some Upper Devonian), and again in the westward promontories north and south of Weston-super-mare. The Somerset Levels form a lowland behind Bridgwater Bay, and to the west the Lower Lias outcrops in low cliffs between Lilstock and Blue Anchor Point. Keuper Marl occupies the Minehead lowland and the valley opening to Porlock, and the Devonian rocks on the northern limb of the mid-Devon syncline form the steep coast at Selworthy Beacon and along the  wnorthern slopes of Exmoor between Porlock and Ilfracombe. The Devonian rocks strike ESE–WNW and the coast cuts successively across Lower and Middle Devonian formations as far as Morte Point and Upper Devonian ­formations south to Saunton. Strongly folded Carbonif­ erous Culm Measures (dark grey and greenish shales and

s­ andstones with grit bands and some limestones) then dominate the coast from the ­Taw-Torridge estuary to Hartland Point and south to the border with Cornwall. Mean spring tide range is large in the Severn estuary, attaining more than 12 m off Avonmouth. As the estuary widens into the Bristol Channel the tide range diminishes from 11.9 m at Clevedon to 11.3 m at Weston-super-mare, 10.5 m at Watchet, 9.9 m at Minehead, 9.5 m in Porlock Bay and 8.5 m at Ilfracombe, and as the Bristol Channel opens to the Atlantic coast they diminish further to 7.4 m at Clovelly 7.2 m on Lundy Island, and just under 7 m at  Hartland Quay. These megatidal conditions result in strong currents in the Severn estuary and wide intertidal zones, notably in Bridgwater Bay where marshes and mudflats are extensive.

1.  Gloucestershire On the Gloucestershire coast of the Severn estuary at Aust there is a cliff cut in Triassic sediment: the Red Marls

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Gloucestershire, Somerset and North Devon

⊡⊡ Fig. 7.7.1.4 Erosion of the Severn-side salt marsh terrace at Tites Point, near Purton. (Courtesy Geostudies.)

References Allen JRL (1992) Tidally induced salt marshes in the Severn estuary, southwest Britain. In: Allen JRL, Pye K (eds) Saltmarshes:

morphodynamics, conservation and engineering significance. Cambridge University Press, Cambridge, New York, pp 123–147 Pethick JS (1996) The geomorphology of mudflats. In: Nordstrom KE, Roman CT (eds) Estuarine shores. Wiley, New York, pp 185–211

7.8  Gloucestershire, Somerset and North Devon The coast of Gloucestershire, Somerset and North Devon begins where the Severn estuary runs beside the Vale of Belvoir, a lowland of Keuper Marl. The Jurassic formations of the Cotswolds do not extend to the coast, but there are minor outcrops of Trias at Aust. The Carboniferous Limestone of the Bristol district reaches the seaboard between Portishead and Clevedon (with some Upper Devonian), and again in the westward promontories north and south of Weston-super-mare. The Somerset Levels form a lowland behind Bridgwater Bay, and to the west the Lower Lias outcrops in low cliffs between Lilstock and Blue Anchor Point. Keuper Marl occupies the Minehead lowland and the valley opening to Porlock, and the Devonian rocks on the northern limb of the mid-Devon syncline form the steep coast at Selworthy Beacon and along the  wnorthern slopes of Exmoor between Porlock and Ilfracombe. The Devonian rocks strike ESE–WNW and the coast cuts successively across Lower and Middle Devonian formations as far as Morte Point and Upper Devonian ­formations south to Saunton. Strongly folded Carbonif­ erous Culm Measures (dark grey and greenish shales and

s­ andstones with grit bands and some limestones) then dominate the coast from the ­Taw-Torridge estuary to Hartland Point and south to the border with Cornwall. Mean spring tide range is large in the Severn estuary, attaining more than 12 m off Avonmouth. As the estuary widens into the Bristol Channel the tide range diminishes from 11.9 m at Clevedon to 11.3 m at Weston-super-mare, 10.5 m at Watchet, 9.9 m at Minehead, 9.5 m in Porlock Bay and 8.5 m at Ilfracombe, and as the Bristol Channel opens to the Atlantic coast they diminish further to 7.4 m at Clovelly 7.2 m on Lundy Island, and just under 7 m at  Hartland Quay. These megatidal conditions result in strong currents in the Severn estuary and wide intertidal zones, notably in Bridgwater Bay where marshes and mudflats are extensive.

1.  Gloucestershire On the Gloucestershire coast of the Severn estuary at Aust there is a cliff cut in Triassic sediment: the Red Marls

Gloucestershire, Somerset and North Devon

­ verlain by Tea Green Marls and a capping of dark grey o shales (>Fig. 7.8.1). To the SW the coast is fringed by salt marshes that have been partly reclaimed behind sea walls. At Avonmouth at low tide, wide, muddy, sandy and gravelly flats are exposed. Low cliffs of Carboniferous Limestone run out to Battery Point and Denny Island is a limestone outcrop offshore, while in Kilkenny Bay salt marshes front the esplanade. Steep coastal slopes in Carboniferous Limestone and Old Red Sandstone descend to scrubby bluffs and low basal cliffs behind rocky and gravelly shores, with mud exposed at low tide. Limestone cliffs continue to the west until River Kenn flows out by Wain Hill. The coast is then low-lying and marshy, with high tide embankments interrupted by the mouth of the River Yeo; then St Thomas’s Head marks the beginning of Middle Hope, a narrow coastal ridge of Carboniferous Limestone mantled by Pleistocene Head deposits. The ridge declines along a promontory to Sand Point, and the steep slopes and low basal cliffs run back along the northern shore of Sand Bay (>Fig. 7.8.2). Sand Bay is a 3 km wide west-facing bay between promontories of Carboniferous Limestone, backed by a sandy beach and low dunes. At low tide up to 2 km of sand and mud is exposed, with numerous sub-parallel creeks running seaward. Near the northern end is a Spartina marsh (>Fig. 7.8.3). Weston-super-mare stands behind Weston Bay, which is similar to Sand Bay, with a wide area of runnelled, silty

⊡⊡ Fig. 7.8.1 Aust Cliff. (Courtesy Geostudies.)

7.8

sand passing seaward below low tide mudflats. It is ­bordered southward by Brean Down, a ridge of Carboniferous Limestone running out to Howe Rock. The River Axe opens into the southern part of the bay, delivering muddy sediment to the shore.

2.  Somerset The long, narrow steep-sided ridge of Carboniferous Limestone running out to Brean Down rises to 97 m, and declines westward to Howe Rock. Steep Holm and  Flat Holm are limestone islands offshore. The Carboniferous Limestone dips steeply north, so that a dip-slope faces Weston Bay and the southern slopes are an escarpment with basal cliffs. To the south, intertidal sand and mudflats are backed by a wide sandy beach and a series of parallel dune ridges. At Burnham on Sea the sandy beach gives way to salt marshes at the mouth of the Brue estuary, which opens to a muddy shore. The larger Parrett estuary comes in from the south beside Stert Point and the elongated Stert Island, a shingle patch which rises above smoothly-sloping mudbanks offshore. There has been mud accretion in Bridgwater Bay and rapid spread of marsh west of the mouth of the Parrett since Spartina was introduced in about 1928 (Ranwell 1964). To the west, the marshes fade out and shingle beaches and beach ridges are arranged in a series of elongated spits

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⊡⊡ Fig. 7.8.2 Limestone promontory north of Sand Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.3 Spartina in Sand Bay. (Courtesy Geostudies.)

with recurves indicating intermittent eastward growth. A  backshore embankment becomes a concrete wall at Stolford, and a gravelly beach that extends westward to Hinkley Point consists of limestone and sandstone pebbles that have drifted from disintegrating Liassic rock outcrops to the west. Hinkley Point power station stands on a low plateau of Liassic limestone and to the west the coastline consists of cliffs fronted by broad shore platforms cut across undulating layers of grey Lias limestone and blue shale, which

­ utcrop in swirling patterns on the wide foreshore o exposed as the tide falls (>Fig. 7.8.4). To the west the vertical cliffs in stratified Lias rise along the seaward side of a coastal ridge to Quantocks Head (over 50 m), past the Blue Ben headland to St Audrie’s Bay. This has a wide foreshore of red sand and sandy mud, backed by disintegrated boulders, a sector of sandy beach at mean high tide, and an upper beach of grey and pink gravel along the base of red cliffs. West of St Audrie’s Bay, the Lias cliffs continue to Watchet. To the west Lias

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7.8

⊡⊡ Fig. 7.8.4 Cliffs and shore platform west from Kilve. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.5 Cliff near Blue Anchor Point cuts across a fault that juxtaposes Lower Lias limestone and shale on the left and Keuper Marl on the right. (Courtesy Geostudies.)

a­ lternates with Keuper Marl in cliff outcrops along Warren Bay to Blue Anchor Point (>Fig. 7.8.5). West from Blue Anchor is a coastal plain fringed by a shingle beach, extending out to a lobate foreland, and then fronting Minehead, where it has been artificially augmented. Beyond Minehead harbour a steep forested slope rises past Culvercliff to Greensleigh Point, and continues along the northern slopes of Selworthy Beacon in steeply dipping, much faulted hard Devonian sandstone.

At Hurlstone Point the steep slope runs in on the eastern side of the Porlock lowland. Porlock Bay has a long beach of cobbles and pebbles (Jennings et al. 1998), backed by a depositional plain that includes areas of reclaimed wetland. The beach was frequently breached by storms, but the breaches generally healed within a week. A breach formed in 1998 has persisted, and the area flooded by high tides has become a salt marsh. Former drains have become wider and deeper tidal creeks, and waterfalls form during the ebb (>Fig. 7.8.6).

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⊡⊡ Fig. 7.8.6 Waterfall over a receding salt marsh cliff, Porlock Bay. (Courtesy Geostudies.)

The beach curves round to the little harbour at Porlock Weir, in the lee of a spit of pebbles and cobbles, the Porlock Weir Ridge, which diverges from the coast to the west and is backed by a grassy marsh. The spit grew as the result of  eastward drifting of beach material derived from cliff recession and slumping and shore erosion of Devonian sandstones to the west, but the supply of pebbles and cobbles appears to have diminished during recent decades, and the spit is unstable. At its western end it passes beneath the steep coast of Culbone (>Fig. 7.8.7) to the Devonshire border at Glenthorne.

3. The North Devon Coast and Lundy Island West from the deep valley at Glenthorne, the bold coast of North Devon curves gradually out to Foreland Point (Arber 1911). Cut in Devonian formations, it is a high, steep hogsback coast with convex slopes curving down to small basal cliffs and a rocky boulder-strewn shore. The steep coast is interrupted at Lynmouth (>Fig. 7.8.8) where the gorges of the East and West Lyn, cut through Devonian slates and grits, converge, and a river descends over ­boulders to a deltaic fan of rock debris exposed at low tide. After wet weather, and when snow melts on the moors, the Lyn discharges large quantities of water, and there was a devastating flood in 1952.

Forested slopes rise along the coast west of Lynmouth, and a high craggy coastal ridge of hard shales and grits is crowned with tors, and has buttresses and screes on the steep coastal slope. The ridge is backed by The Valley of Rocks, a dry valley running parallel to the coast and opening through a gap to Lee Bay. At Heddon’s Mouth, a deep V-shaped valley has been cut across the strike of the Lower Devonian rocks, and the steep coastal slopes continue westward to Blackstone Point, where the slopes rise abruptly to Great Hangman (318 m) and Little Hangman (218 m). Combe Martin is a long, straight, incised valley, and at its mouth is a beach of pebbly sand between planed-off outcrops of steeply dipping sandstone. To the west a coastal ridge on southward dipping slates is divided by inset sandy coves at Sandy Bay and Small Mouth, and backed by a long inlet, Watermouth, which drains out at low tide to expose sand and gravel. The coast then rises and steepens to Widmouth Head and Rillage Point. The deep Chambercombe valley opens to Hele Bay, and the steep coast then runs out to Beacon Point above a rocky shore. On its western side a series of coves and rocky ribs extends to Larkstone Beach and the bay that is on its western side has a sandy beach in Ilfracombe Harbour, sheltered by a breakwater. Ilfracombe occupies a steep-sided east-west valley behind a coastal ridge interrupted by sandy Wildersmouth Bay. To the west, a steep coast runs along Seven Hills,

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7.8

⊡⊡ Fig. 7.8.7 The western end of the Porlock Weir Ridge, showing the steep coast towards Culbone. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.8 The steep coast at Lynmouth, showing a lowtide delta of boulders and gravel carried down by the River Lyn during episodes of flooding. (Courtesy Geostudies.)

interrupted by a deep valley descending to Lee Bay. Morte Point runs out sharply westward with scrubby knolls that become craggy and projecting slate tors like teeth on a dragon’s back. The coastal slopes are mantled with periglacial rubble (Head) and descend to cliffs of ­collapsing grey slate that overlook a white water race as tides meet waves. Lundy Island is about 30 km west of Morte Point. It is almost flat-topped, rising 100–120 m above sea level, and

consists of granite, with a small outcrop of Devonian rock at the SE end. Its coasts show a slope-over-wall profile, the slope mantled by Pleistocene periglacial Head deposits above granite cliffs that are generally higher on the more exposed western side. Morte Bay has been cut out in relatively soft shales, and contains Barricane Beach, famous for shell accumulations in clefts between rocky ribs. Woolacombe Sand has surf breaking on fine, firm, wide, calcareous sands

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exposed at low tide, with megaripples and pools emerging as the tide ebbs. At the southern end is the broad Baggy Point peninsula with steep rock faces, and the craggy cliffs run back into sandy Croyde Bay, at the mouth of a valley cut into southward-dipping Upper Devonian beds. At Saunton Head-mantled slopes descend to an emerged shore platform with outcrops of bedded calcareous sandstone, a Pleistocene beach and dune sequence that incorporates glacial erratics (Campbell and Gilbert 1998). Saunton Sands are up to 500 m wide at low tide, backed by dunes (Braunton Burrows) up to 30 m high. At their southern end a recurved spit - Crow Point - borders the wide TawTorridge estuary where at low tide the river channel flows to the sea between wide sands, with waves breaking over Bideford Bar (Kidson 1963). At Fremington Quay on the southern side of the Taw estuary the river at low tide meanders close to the shore. To the east a beach fringes low cliffs cut in earthy gravel containing rounded and angular pebbles, mainly of culm but also some quartzite, basalt, Old Red Sandstone, Scottish erratics (including dolerite, granite, quartzite, andesite, tuff, gneiss and flint) as well as Irish Sea foraminifera. These erratics indicate a Pleistocene glacial drift deposit. Appledore is a small port on a promontory at the mouth of the Torridge River, where mean spring tide range is 7.4 m. The coast is walled, but there is an intertidal area of angular culm gravel and sand shoals.

Dissected salt marsh exists in a bay to the west, behind a broad foreland, Northam Burrows. On the western shore, extending north from Westward Ho!, is the Pebble or Popple Ridge, a barrier spit 25 m wide and up to 3.5 m above high spring tide level, consisting of rounded boulders (>Fig. 7.8.9). From Westward Ho! cliffs cut in the Culm Measures and truncate a series of east-west strike ridges and valleys. Steep scrubby coastal slopes are fronted by beaches of grey and pink gravel drifting east towards the Popple Ridge at Westward Ho! At Portledge, cliffs are cut in an outlier of Triassic Keuper Marl with sandstone and breccia layers (>Fig. 7.8.10). At Peppercombe the steep coast descends to a beach of coarse grey cobbles and boulders. East of the mouth of Peppercombe valley the steep forested coastal slopes have hollows and benches aligned with the geological strike, formed by slumping. Buck’s Mills is in a deeply incised forested valley which interrupts the steep coastal  escarpment. Clovelly stands on an anticlinal pro­montory, and the coast westward to Hartland Point is steep and wooded, incised by deep valleys, as at Mouth Mill. At Hartland Point the coast turns southward to face the Atlantic. High vertical cliffs rise to a plateau 60–80 m above sea level that truncates the strongly folded Culm Measures exposed in the cliffs and on the shore. There are smooth steeply inclined bedding planes in the cliffs and truncated strata and rocky ribs on the shore. ⊡⊡ Fig. 7.8.9 The Popple Ridge at Northam Burrows. (Courtesy Geostudies.)

Gloucestershire, Somerset and North Devon

Several short, steep streams descend from a drainage divide a few kilometres inland to end in coastal waterfalls (>Fig. 7.8.11). Caves, ribs and pillars have been cut out in sharp anticlines north of Hartland Quay, where the cliffs rise 100 m and the cliff base is fringed by intertidal shore platforms, where fan-folded strata have been dissected.

⊡⊡ Fig. 7.8.10 Cliffs of Keuper Marl at Portledge. (Courtesy Geostudies.)

⊡⊡ Fig. 7.8.11 Waterfall at Speke’s Mill Mouth. (Courtesy Geostudies.)

7.8

Between Speke’s Mill Mouth and Welcombe the cliffs are more than 120 m high. Several buttress reefs of thick hard shale run out from the coast, forming an irregular intertidal shore; Knaps Longpeak (>Fig. 7.8.12) is a salient of hard shale. At Welcombe Mouth a deep valley with ­bordering terraces of gravelly Head opens between cliffs to a rocky shore. A steep bluff extends to the mouth of the

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The North Coast of Cornwall

⊡⊡ Fig. 7.8.12 Knaps Longpeak, on Carboniferous (Culm sand­­stone. (Courtesy Geostudies.)

steep-sided Marsland valley, which marks the boundary between Devon and Cornwall.

References Arber EAN (1911) The Coast Scenery of North Devon. Dent, London Campbell S, Gilbert A (1998) The Croyde-Saunton coast. Quaternary of South-West England. Geol Conserv Rev Ser 14:214–224

Jennings SC, Orford JD, Canti M, Devoy RJN, Straker V (1998) The role of relative sea-level rise and changing sediment supply on Holocene gravel barrier development: the example of Porlock, Somerset. Holocene 8:165–181 Kidson C (1963) The growth of sand and shingle spits across estuaries. Zeitschrift für Geomorphologie 7:1–22 Ranwell DS (1964) Spartina salt marshes in southern England: Bridgwater Bay. J Ecol 52:79–94

7.9.1  The North Coast of Cornwall 1.  Introduction The north coast of Cornwall is bold, and receives strong swell which has been transmitted from stormy areas in the Atlantic Ocean, as well as occasional large waves generated by gales in the Western Approaches (Bird 1998). Mean spring tide ranges are from 5.9 m at St. Ives and 6.2  m at Perranporth to 6.6 m in the Camel estuary at Padstow. The coast is relatively straight, with high cliffs and rocky headlands, and short steep valleys cut by streams, some of which end in coastal waterfalls. There are  slope-over-wall coasts with steep, vegetated slopes descending to basal rocky cliffs, particularly on the flanks of headlands, and in several places the coastal slope has been removed by erosion and there are vertical cliffs, as in Watergate Bay north of Newquay and on the coast north

of Bude. The highest point on the coast is 213 m at the crest of the steep slope at High Cliff, north of Tintagel, but the highest rocky cliffs are between Navax Point and St. Agnes Head, from Newquay to Trevose Head, on Pentire Point and north of Bude.

2.  The Coastline South of the Devonshire border at Marsland Mouth the coast rises to steep cliffs cut in strongly folded and much faulted Carboniferous Culm Measures (Bude Formation). The Bude Formation consists of sandstones and shales in outcrop in sharp anticlines and synclines exposed and dissected in the cliffs and shore platforms up to 200 m wide at low spring tides. A dissected plateau comes to the coast,

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The North Coast of Cornwall

⊡⊡ Fig. 7.8.12 Knaps Longpeak, on Carboniferous (Culm sand­­stone. (Courtesy Geostudies.)

steep-sided Marsland valley, which marks the boundary between Devon and Cornwall.

References Arber EAN (1911) The Coast Scenery of North Devon. Dent, London Campbell S, Gilbert A (1998) The Croyde-Saunton coast. Quaternary of South-West England. Geol Conserv Rev Ser 14:214–224

Jennings SC, Orford JD, Canti M, Devoy RJN, Straker V (1998) The role of relative sea-level rise and changing sediment supply on Holocene gravel barrier development: the example of Porlock, Somerset. Holocene 8:165–181 Kidson C (1963) The growth of sand and shingle spits across estuaries. Zeitschrift für Geomorphologie 7:1–22 Ranwell DS (1964) Spartina salt marshes in southern England: Bridgwater Bay. J Ecol 52:79–94

7.9.1  The North Coast of Cornwall 1.  Introduction The north coast of Cornwall is bold, and receives strong swell which has been transmitted from stormy areas in the Atlantic Ocean, as well as occasional large waves generated by gales in the Western Approaches (Bird 1998). Mean spring tide ranges are from 5.9 m at St. Ives and 6.2  m at Perranporth to 6.6 m in the Camel estuary at Padstow. The coast is relatively straight, with high cliffs and rocky headlands, and short steep valleys cut by streams, some of which end in coastal waterfalls. There are  slope-over-wall coasts with steep, vegetated slopes descending to basal rocky cliffs, particularly on the flanks of headlands, and in several places the coastal slope has been removed by erosion and there are vertical cliffs, as in Watergate Bay north of Newquay and on the coast north

of Bude. The highest point on the coast is 213 m at the crest of the steep slope at High Cliff, north of Tintagel, but the highest rocky cliffs are between Navax Point and St. Agnes Head, from Newquay to Trevose Head, on Pentire Point and north of Bude.

2.  The Coastline South of the Devonshire border at Marsland Mouth the coast rises to steep cliffs cut in strongly folded and much faulted Carboniferous Culm Measures (Bude Formation). The Bude Formation consists of sandstones and shales in outcrop in sharp anticlines and synclines exposed and dissected in the cliffs and shore platforms up to 200 m wide at low spring tides. A dissected plateau comes to the coast,

The North Coast of Cornwall

ending in steep slopes down to rugged cliff and shore scenery, with high headlands and deep valleys (cleaves). The incised Yeolmouth valley has a waterfall pouring  down into a shingle cove, and near Morwenstow, another deep V-shaped valley truncated by cliff recession ends in a waterfall above St. Morwenna’s Well. There is a recent major landslide south of Higher Sharpnose Point (>Fig. 7.9.1.1). To the south a deeply incised valley opens to Stanbury Mouth, which has a cliff waterfall and a beach of sandy and grey gravels between rocky ribs. South from Warren Point the proportion of rocky ribs and shores diminishes as sandy beaches lengthen towards Bude. At Earthquake a major landslide has occurred down a sandstone bedding plane, producing a basal fan of talus, and large sandstone ribs run out into sea. Bude is built behind a sandy bay at the mouth of the River Neet valley. The beach sands are calcareous, including shelly debris washed in from the sea floor and quartzose sand derived from the erosion of weathering sandstones exposed in the nearby cliffs and shore platforms. Large quantities of shellysand were extracted from this beach in the eighteenth and nineteenth centuries, resulting in the reduction of Summerleaze Beach. Whale Back (folded sandstones) and Saddle Rock are en echelon pericline folds, which look like inverted boats when seen from the top of the cliff. To the south, marine erosion has cut steep to vertical cliffs across the sharply folded sandstones, mudstones and shales, fronted by a gently sloping intertidal shore platform cut across these formations, ⊡⊡ Fig. 7.9.1.1 The steep Atlantic coast south of Higher Sharpnose Point. (Courtesy Geostudies.)

7. 9.    1

truncating the anticlines and synclines that run at right angles to the coastline. Variations in the resistance of the outcropping strata are expressed as ribs and clefts in the cliffs and ridges and furrows across the shore platform (>Fig. 7.9.1.2). The promontory of Lower Longbeak, cut in massive mudstones, marks the southern end of the high cliffs. Widemouth Bay is backed by low cliffs cut in Head Deposits. The shore platforms show structural features developed across strongly-folded sandstone and shale formations in which the axes of folding run out westward across the shore. The sandstones and shales of the Bude For­mation give place to dark shales and thin grey sandstones of the Crackington Formation at Wanson Mouth. This formation has some outcrops mainly of sandstone, forming bold cliffs and resistant outcrops on the shore, and others (generally narrow) mainly of shale, excavated to form coves, inlets and caves. It dominates the coast south to Boscastle. Much of the coast is steep, with slope-overwall profiles cut into fault-bounded outcrops of various Upper Devonian and Lower Carboniferous formations. At Crackington Haven two deep valleys meet behind a sandy cove, backed by grey shingle and low cliffs in Head, between cliffs and steep slopes. As the coast runs out to Cambeak the cliffs are cut across strongly folded sandstones and shales. Cambeak (98 m) is a sloping sandstone headland with tumbled grassy slopes in Head and cliffs that show swirling folds and recumbent folding in the Crackington Formation. South of Fire Beacon Point there are landslides on steep coastal slopes to Beeny Cliff

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⊡⊡ Fig. 7.9.1.2 Shore platform cut across steeply dipping sandstones and shales of the Bude Formation at Lower Longbeak, south of Bude. (Courtesy Geostudies.)

(>Fig. 7.9.1.3), where convex coastal slopes intersect northward-dipping sandstones and descend to rocky cliffs broken by large caves. South towards Boscastle high ground recedes from the coast and the Trevena Platform (90–130 m) dominates the hinterland, ending in bevelled slope-over-wall profiles. Boscastle Harbour is the drowned mouth of the incised Valency valley, bordered by steep craggy slopes in the strongly folded slates and sandstones of the Crackington Formation. The coast between Boscastle and Bossiney Haven has many inlets at the mouths of deeply-incised valleys where streams flow steeply down to the coast. South of Trevalga Head a deep cove has been cut in rocks that dip steeply northeastward. Cliffs in green Tredorn Slates are cut by numerous low-angle faults and some steeper, almost vertical fractures. The Tintagel Volcanic Formation is prominent in the high craggy Trevalga Cliffs and Rocky Valley is a long and deep valley cut in Delabole Slates that descends to the coast. There is a sandy beach at low tide in Bossiney Haven. In the Tintagel district cliff outlines have been influenced by dissection along rectilinear joint patterns in Devonian and Carboniferous slates and volcanic rocks (Wilson 1952). Upper Devonian slates cap Tintagel Island, thrust up over Carboniferous Tintagel Volcanics. South of Tintagel cliffs, Devonian and Carboniferous rocks extend along Port Isaac Bay, past Port Quin, a curving drowned valley-mouth inlet, and out to the bold promontory between Rumps Point and Pentire Point. This borders the wide Camel estuary, which has extensive

inwashed sands. On its northern shore is the Trebetherick raised beach (>Fig. 7.9.1.4), on an emerged shore platform rising to about 3 m above high tide level, cut across the dipping slates (Arkell 1943). This dates from a phase when the sea stood higher than it does now, and shows features that formed during an ensuing phase of cold (periglacial) climate and much lower sea level. The south coast of the Camel estuary shows a succession of promontories on hard igneous rocks, notably greenstones, lava and tuff, and intervening bays cut out in the softer slates. Between Stepper Point and Trevose Head is a broad bay backed by low cliffs and coastal slopes of subdued outline, cut in weathered Middle to Upper Devonian slates. South of Porthmissen Bridge the cliff crest curves round the head of a deep bouldery inlet, the southern side of which is Marble Cliff, over 30 m high, cut in about 140 thin bands of alternating gently-dipping grey limestone, dark shale and striped grey slate. This is the lower part of an overfold, so that the strata are inverted, the youngest beds being near the bottom. Towards Trevose Head the steep coastal slope is cut in massive dark greenstone. Trevose Head and Dinas Head are windswept headlands, and Constantine Bay is a gentlycurving west-facing bay cut in shales, and occupied by a wide surf beach of shelly sand, backed by dunes held in place by marram grass. To the south the cliffs have been much dissected by marine erosion, isolating numerous stacks and small islands. At Bedruthan Steps (>Fig. 7.9.1.5) the cliffs in Devonian slates are almost vertical, and there are several high stacks including the pinnacle of Queen Bess Rock.

The North Coast of Cornwall

7. 9.    1

⊡⊡ Fig. 7.9.1.3 The slope-over-wall profile of Beeny Cliff, north of Boscastle, with High Cliff in the ­background. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.4 The emerged beach at Trebetherick Point, on the eastern shores of the Camel estuary, resting on a bench cut into Upper Devonian slates and overlain by laminated dune calcarenite. Brae Hill in the distance is an outcrop of greenstone. (Courtesy Geostudies.)

On Trenance Point there are sandstones and shales with hard quartzite forming protrusions on the bold cliffs. Mawgan Porth is a sanded inlet at the mouth of the Vale of Mawgan. High cliffs border Beacon Cove, and at Strasse Cliff the Watergate Bay surf beach begins. The coastline of Watergate Bay is cut across an E–W  anticline, where the Meadfoot Beds overlie Dart­­ mouth Beds rising through the cliffs. It has a sandy beach 2.5 miles long in front of vertical cliffs at the edge of the St. Mawgan plateau. Ocean swell washes the base

of  the cliffs at high tide, and rip currents are powerful (>Fig. 7.9.1.6). At Whipsiderry the coast turns westward, and the Meadfoot Beds continue along the north coast of Porth Island, where the southward dip produces an escarpment cliff. The great cave known as the Banqueting Hall was formerly about 60 m long, up to 20 m wide and nearly 20 m high, with two entrances from the sea and a third through the roof. Unfortunately it became unsafe and the local council felt obliged to destroy it with explosives

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⊡⊡ Fig. 7.9.1.5 The cliffs, stacks and low-tide sandy beach near Bedruthan Steps, looking towards Park Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.6 An air view of Watergate Bay, with strong Atlantic swell at high tide. The receding cliffs retain no remnants of the ­periglaciated slope seen on less exposed parts of the Cornish coast. (Courtesy Geostudies.)

in 1987. Porth Island runs out to Trevelgue Head (>Fig. 7.9.1.7), and is bordered by the long inlet at Porth, where sand is exposed at low tide. The Newquay coast has cliffs fronted by flat brown sands exposed at low tide. Towan Head is a rounded promontory of hard calcareous slates and thin limestones running out to the northwest. A low cliff exposes Holocene dune sand over gravelly Head up to 3 m of thinly bedded Pleistocene calcareous sandrock (a beach or dune calcarenite), with an emerged beach resting upon grey Mead­foot Slates.

Fistral Bay is cut into rapidly weathering shales, and has a sandy beach washed by surf from Atlantic swell, backed by grassy dunes. Pentire Point East and West are headlands in hard slates and grits, divided by The Gannel, another sanded inlet. Porth Joke, to the west, is similar. South of Kelsey Head is Holywell Bay, several caves below calving cliffs capped by Head and cliff-top dunes, and a wide beach backed by grassy dunes. Cliffed headlands run out to Penhale Point and Ligger Point, then Perran Sands, backed by high dune topography that extends just over 2 miles south to Perranporth.

The North Coast of Cornwall

7. 9.    1

⊡⊡ Fig. 7.9.1.7 An air view of Trevelgue Head and Porth at high tide, with the cove at Lusty Glaze on the right. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.8 The stack at Tobban Horse, a stack south of Porthtowan, with the remains of the coastal slope on its seaward side. Erosion during the past 6,000 years has excavated a cove here, cutting the cliffs back by up to 200 m. (Courtesy Geostudies.)

Droskyn Point was much modified by mining and slate quarrying in the eighteenth and nineteenth centuries, and a rugged slope-over-wall coast runs out to Cligga Head and on to St. Agnes Head and Navax Point. The proportion of slope to wall varies according to exposure to Atlantic waves. South of Porthtowan is Tobban Horse, a large stack with a grassy Head-mantled slope on its seaward side, from which the extent of cliff recession can be estimated (>Fig. 7.9.1.8).

At Portreath is a valley-mouth beach of calcareous sand mixed with sand and gravel from mining waste. To the SW, even-crested cliffs about 80 m high truncate the coastal plateau (Trevena Platform) and a slope-over wall coast runs out to Naxax Point. Godrevy Point marks the beginning of St. Ives Bay, and the cliffs cut in slates at Godrevy include an emerged Pleistocene beach and stratified dune calcarenite beneath Head deposits, as well as a fronting shore platform (>Fig. 7.9.1.9).

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⊡⊡ Fig. 7.9.1.9 Cliffs cut in Head deposits back the broad shore platform at Godrevy, on the north coast of St. Ives Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.1.10 Granite tor above Gamper, Land’s End, looking along the slope-over-wall coast to Pedn-men-du, with the stack known as the Irish Lady (arrowed). (Courtesy Geostudies.)

Red River comes down to the sea from the mines of the Camborne district. It was formerly stained by oxidised waste from the mining area, but this pollution has been halted. The beach sands of St. Ives Bay contain tin and other heavy minerals supplied by streams draining from mineral-rich hinterlands and eroded from coastal outcrops. They are backed by dunes that have been drifting inland since mediaeval times, burying farmland and early settlements.

The River Hayle flows through a broad estuary to reach the sea through an outlet between the coastal dunes behind Porth Kidney Sands. To the west cliffs and bluffs extend to St. Ives. The north coast of the Land’s End (Penwith) Peninsula has sectors of granite coast between extensive fringes of metamorphic rock and greenstone intrusions. It is a high, steep and cliffy coast facing stormy seas, and its outlines have been strongly influenced by patterns of jointing in the rocks. Zennor Head is a promontory of

The South Coast of Cornwall

greenstone (dolerite) with craggy slopes at the edge of a plateau about 90 m above sea level. Typical granite cliffs, dissected along vertical and almost horizontal joint planes, run from Gurnards Head to Portheras Cove, where there is a junction between the granite and Mylor Slates. The cliffs are incised by deep zawns, steep-sided inlets cut out along joints or fault planes, some of which have been excavated or enlarged by mining along tin and copper lodes in  the coastal rock outcrops. Beyond Cape Cornwall, in Priest’s Cove the metamorphic slates give place to granite, which dominates the coast from here to Lands End. At Porth Nanven there is an emerged beach of granite boulders, overlain by gravelly Head deposits. Whitesand Bay has an upper beach of coarse shelly brown sand behind a wide almost flat area of finer softer mica-rich sand exposed at low tide; to the south, Sennen, Cove has a beach of cobbles and boulders.

7. 9.       2            

The coast curves south, past the cliffed headland of Pedn-men-du and the Irish Lady, a stack with a conical hat. Grassy slopes on Head deposits have granite tors (>Fig. 7.9.1.10) and are trenched by large cliff cauldrons, possibly collapsed caves, as the coast runs out to the castellated granite cliffs of Dr. Syntax’s Head, the promontory that declines to Land’s End.

References Arkell WJ (1943) The Pleistocene rocks at Trebetherick Point, Cornwall: their interpretation and correlation. Proc Geol Assoc 54:141–170 Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Wilson G (1952) The influence of rock structures on coast-line and cliff development around Tintagel. North Cornwall. Proc Geol Assoc 63:20–49

7.9.2  The South Coast of Cornwall 1.  Introduction The south coast of Cornwall is more subdued and indented than the north coast, and receives waves generated by winds across the English Channel as well as south-westerly ocean swell from the Atlantic (Bird 1998). Wave action is generally light to moderate, with large storm waves arriving only occasionally. Mean spring tide range averages 4.7–4.8 m. Slope-over-wall profiles have steep vegetated slopes descending to rocky cliffs that are usually less than 10 m high (Arber 1949). The proportion of rocky cliff to coastal slope is less than on much of the north coast, but on the western side of the Lizard Peninsula, exposed to Atlan­­ tic  swell and storm waves, vertical cliffs rise locally to 50–70 m. Head­lands and promontories correspond with higher ground or harder outcrops, the intervening bays and coves occupying lower ground or softer outcrops. The coastline runs in along the sides of several large branching inlets known as rias, formed where the sea has invaded the lower parts of river valleys, notably the Tamar, the Helston and Fowey Rivers, and Carrick Roads. From time to time large waves break upon the southern shores of Cornwall. Some are generated by storms, but there are also records of sudden bursts of wave activity in calm weather, particularly in Mounts Bay and along the south coast to Whitsand Bay. Such waves were observed along the south coast on 1 November 1755, 4 h and 15 min after the catastrophic Lisbon earthquake generated a tsunami transmitted northward in the West Atlantic: large

granite boulders were thrown up well above high tide level at Lamorna Cove on the Land’s End Peninsula. Mean spring tide ranges show little variation: at Newlyn 4.8 m, Porthleven, Lizard Point, Coverack, Helford River mouth, Falmouth and Mevagissey 4.7 m, Fowey and Looe 4.8 m and Whitsand Bay 4.5 m.

2.  The Coastline The columnar granite cliffs of Land’s End are dissected along horizontal and vertical joint planes, and interrupted by grassy chutes of Head deposits. Many granite blocks have fallen to the shore. There are valley-mouth coves with little sandy beaches. In places there is a narrow cliff-base ramp, cut out along joint planes inclined gently seaward. The sea is green over a sandy floor, but darker over weedy areas amid inshore boulders. In Pendower Cove there are caves cut out along joint planes (>Fig. 7.9.2.1). The coast to Gwennap Head is bold, ­facing SW into an often stormy sea; even in calm weather Atlantic swell breaks as a heavy surf over reefs and shore rocks. On Gwennap Head rectilinear jointing is ­prominent, and the coast turns eastward. The southern coast of the Land’s End peninsula is almost entirely medium to coarse grained granite, with associated igneous rocks, such as the greenstones at Tater-du. Porthcurno Cove, at the mouth of a steep-sided valley, has a sandy beach backed by dunes, and to the east the coast curves out to the promontory of

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The South Coast of Cornwall

greenstone (dolerite) with craggy slopes at the edge of a plateau about 90 m above sea level. Typical granite cliffs, dissected along vertical and almost horizontal joint planes, run from Gurnards Head to Portheras Cove, where there is a junction between the granite and Mylor Slates. The cliffs are incised by deep zawns, steep-sided inlets cut out along joints or fault planes, some of which have been excavated or enlarged by mining along tin and copper lodes in  the coastal rock outcrops. Beyond Cape Cornwall, in Priest’s Cove the metamorphic slates give place to granite, which dominates the coast from here to Lands End. At Porth Nanven there is an emerged beach of granite boulders, overlain by gravelly Head deposits. Whitesand Bay has an upper beach of coarse shelly brown sand behind a wide almost flat area of finer softer mica-rich sand exposed at low tide; to the south, Sennen, Cove has a beach of cobbles and boulders.

7. 9.       2            

The coast curves south, past the cliffed headland of Pedn-men-du and the Irish Lady, a stack with a conical hat. Grassy slopes on Head deposits have granite tors (>Fig. 7.9.1.10) and are trenched by large cliff cauldrons, possibly collapsed caves, as the coast runs out to the castellated granite cliffs of Dr. Syntax’s Head, the promontory that declines to Land’s End.

References Arkell WJ (1943) The Pleistocene rocks at Trebetherick Point, Cornwall: their interpretation and correlation. Proc Geol Assoc 54:141–170 Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Wilson G (1952) The influence of rock structures on coast-line and cliff development around Tintagel. North Cornwall. Proc Geol Assoc 63:20–49

7.9.2  The South Coast of Cornwall 1.  Introduction The south coast of Cornwall is more subdued and indented than the north coast, and receives waves generated by winds across the English Channel as well as south-westerly ocean swell from the Atlantic (Bird 1998). Wave action is generally light to moderate, with large storm waves arriving only occasionally. Mean spring tide range averages 4.7–4.8 m. Slope-over-wall profiles have steep vegetated slopes descending to rocky cliffs that are usually less than 10 m high (Arber 1949). The proportion of rocky cliff to coastal slope is less than on much of the north coast, but on the western side of the Lizard Peninsula, exposed to Atlan­­ tic  swell and storm waves, vertical cliffs rise locally to 50–70 m. Head­lands and promontories correspond with higher ground or harder outcrops, the intervening bays and coves occupying lower ground or softer outcrops. The coastline runs in along the sides of several large branching inlets known as rias, formed where the sea has invaded the lower parts of river valleys, notably the Tamar, the Helston and Fowey Rivers, and Carrick Roads. From time to time large waves break upon the southern shores of Cornwall. Some are generated by storms, but there are also records of sudden bursts of wave activity in calm weather, particularly in Mounts Bay and along the south coast to Whitsand Bay. Such waves were observed along the south coast on 1 November 1755, 4 h and 15 min after the catastrophic Lisbon earthquake generated a tsunami transmitted northward in the West Atlantic: large

granite boulders were thrown up well above high tide level at Lamorna Cove on the Land’s End Peninsula. Mean spring tide ranges show little variation: at Newlyn 4.8 m, Porthleven, Lizard Point, Coverack, Helford River mouth, Falmouth and Mevagissey 4.7 m, Fowey and Looe 4.8 m and Whitsand Bay 4.5 m.

2.  The Coastline The columnar granite cliffs of Land’s End are dissected along horizontal and vertical joint planes, and interrupted by grassy chutes of Head deposits. Many granite blocks have fallen to the shore. There are valley-mouth coves with little sandy beaches. In places there is a narrow cliff-base ramp, cut out along joint planes inclined gently seaward. The sea is green over a sandy floor, but darker over weedy areas amid inshore boulders. In Pendower Cove there are caves cut out along joint planes (>Fig. 7.9.2.1). The coast to Gwennap Head is bold, ­facing SW into an often stormy sea; even in calm weather Atlantic swell breaks as a heavy surf over reefs and shore rocks. On Gwennap Head rectilinear jointing is ­prominent, and the coast turns eastward. The southern coast of the Land’s End peninsula is almost entirely medium to coarse grained granite, with associated igneous rocks, such as the greenstones at Tater-du. Porthcurno Cove, at the mouth of a steep-sided valley, has a sandy beach backed by dunes, and to the east the coast curves out to the promontory of

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⊡⊡ Fig. 7.9.2.1 Caves on a granite coast in Pendower Cove, near Land’s End. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.2 Logan Rock (L), perched on a cliff-top tor near Treen, on the south coast of the Land’s End Peninsula. The coarsegrained granite is divided by three intersecting joints, two almost vertical planes and one roughly horizontal and curving. (Courtesy Geostudies.)

Horrace, on which is an Iron Age fort and Logan Rock, a 65 tonne rocking boulder (> Fig. 7.9.2.2). The granite cliffs run eastward to Lamorna Cove, where the valley-side slopes are strewn with granite boulders. At Mousehole breakwaters enclose a former cove as a harbour, dry and sandy at low tide. Just south is the ­junction between the granite and the Mylor Slates, here metamorphosed into hard green splintery rock. Low cliffs in these slates extend northward past Penlee Point, an outcrop of greenstone. On South Pier at Newlyn is the tidal observatory, where the

rise and fall of the tide is recorded. There is evidence that mean sea level has been rising here in recent decades. From Newlyn the coast curves NE and is largely artificial, round to Penzance. Mounts Bay has brown muddy sand with outlying rocky reefs and a submerged forest. Behind the sandy beach towards Marazion, grassy dunes are backed by reedswamp and scrub (Marazion Marsh) in a former valley-mouth lagoon. The Causeway (> Fig. 7.9.2.3) that leads out at low tide to St. Michael’s Mount is an ebb road, passable for an hour or two before and after low

The South Coast of Cornwall

tide: the duration of emergence has diminished as the result of sea level rise (as indicated on the Newlyn tide gauge). St. Michael’s Mount is a high island of granite with an outcrop of Mylor Slates on its eastern shore. Granite ledges, sloping gently SE along joint planes, rise along the  southern shore, and granite boulders on the NE shore  were emplaced by storm surges or tsunamis. Low vertical cliffs in Head deposits and reefs and headlands of greenstone extend east from Marazion to Cudden Point. Prah Sands occupy an embayment backed by cliffs in Head ⊡⊡ Fig. 7.9.2.3 The causeway (ebb road) to St. Michael’s Mount crosses a rocky, sandy and bouldery isthmus exposed at low tide. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.4 A fault exposed in the cliffs behind Rinsey Cove, with a cave excavated along the fault plane. (Courtesy Geostudies.)

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and Mylor Slates and in Rinsey Cove the cliff exposes the contact between the Mylor Slates and the Tregonning granite (> Fig. 7.9.2.4). On the shore platform fronting cliffs west of Porthleven is Giant’s Rock (> Fig. 7.9.2.5), a large rounded block of smooth pink microcline gneiss, a rock type unknown in Britain is seen and probably brought here in an iceberg that floated on to the shore and melted. Porthleven Harbour occupies a valley-mouth inlet and to the east the beach consists of fine flinty shingle and

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coarse sand which extends SE as Loe Bar, a barrier ­enclosing Loe Pool in the drowned valley mouths of the River Cober and Carminowe Creek. The beach continues SE to Gunwalloe, where a Late Pleistocene emerged beach can be seen. To the south cliffs extend past headlands and coves to Church Cove, where a headland is capped by a stabilised dune. There is a contrast between the rocks on either side of the sandy beach in Polurrian Cove: Devonian Portscatho slates on the north and hornblende schists to the south.

The schists run round to Mullion Cove, where there are dykes and lenses of serpentine, which dominates the high cliffs south past Kynance Cove towards Lizard Head. Steep and rugged cliffs in mica schists extend along the coast to Lizard Point. The cliffs decline to Housel Bay from Lizard Point where The Lion’s Den (> Fig. 7.9.2.6) is an almost circular hole nearly 50 m in diameter at the top of a shaft over 70  m deep, formed when the roof of a cave collapsed in 1847. ⊡⊡ Fig. 7.9.2.5 Giant’s Rock, a large boulder of gneiss thought to have arrived on the shore west of Porthleven as an ice-borne erratic in Pleistocene times. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.6 Slope-over-wall cliffs on the western side of Housel Bay, Lizard Peninsula, showing cliff cauldrons and the Lion’s Den (L), a round hole formed by the collapse of a cave roof. (Courtesy Geostudies.)

The South Coast of Cornwall

Steep rugged cliffs with crags and tors run out on Pen Olver, where the coast turns northward. Steep grassy slopes descend to low cliffs in hornblende schist along this more sheltered coast to Cadgwith and Kennack Sands, and then higher cliffs run eastward to Black Head. At Coverack the shore is irregular, with serpentine outcrops cut by dykes of coarse crystalline gabbro, black epidiorite and other dark intrusive rocks. To the north on Lowland Point a wide terrace is cut by low cliffs in Head deposits. Godrevy Cove, Porthoustock and Porthallow are bays on a cliffed coast cut in gabbro, with beaches ­augmented by gravelly quarry waste. Beyond Nare Point the cliffs recede and decline beside drowned valleys at Gillan Harbour and the Helford River. The rounded peninsula of Rosemullion Head is cut into the dark contorted slates and sandstones of the Portscatho Beds, and to the north the cliffs rise to a succession of rocky promontories and small sandy coves, with ridges of slate running out eastward across the shore. Maen­porth has a sandy beach behind a wide valley-mouth cove, and near Newporth Head there are prominent cliff-base ledges bearing remnants of a gravelly raised beach (> Fig. 7.9.2.7). The coast at Falmouth consisted of receding cliffs of Head deposits stabilised by construction of a sea wall. This is fronted by a gravelly beach and a shore platform that is being cut into a similar Late Pleistocene shore platform, which stands 2–3 m higher. Pendennis Point is fringed by vegetated slopes that descend to cliffs of rubbly Head above steep rocky shores. This marks the entrance to

⊡⊡ Fig. 7.9.2.7 Sunny Cove, near Falmouth, showing remnants of the Late Pleistocene emerged shore platform (P) between segments of modern shore platform backed by cliffs. (Courtesy Geostudies.)

7. 9.       2            

Carrick Roads, a large branching drowned valley system. The River Fall, which opens into an upper branch of Carrick Roads, built a delta of china clay brought down from quarries on Hensbarrow Down (Ranwell 1974), but when this clay pollution was halted the delta began to erode (> Fig. 7.9.2.8). St. Mawes stands behind a south-facing bay on a peninsula north of the Porthcuel River. The shores are rocky, cut into outcrops of the soft slates of the Portscatho Formation. At St. Anthony Head, on the eastern side of the entrance to Carrick Roads, the coast swings round to the NE. The cliffs and shore rocks are interrupted by Towan Beach, mainly sandy with some gravel, in front of low cliffs of Head and a section of coastal dunes. North of Portscatho the cliffs follow smooth cleavage planes and the coastline runs along the strike of the shore outcrops. The shore becomes sandy as the coast curves behind Gerrans Bay, and exposure to SW waves increases. There are emerged shore platforms and segments of raised beach, then the coast curves south towards Nare Head. From Nare Head to Portholland the cliffs of Veryan Bay are cut mainly in the Roseland Breccia, a jumbled mass of brecciated slates and sandstones. There are sand and gravel beaches at Caerhayes and Hemmick, and the Dodman is a cliffed promontory of phyllites - the cliffs higher on the more exposed western coast (> Fig. 7.9.2.9). Vault Beach (> Fig. 7.9.2.10), to the east, is sandy, with patches of white quartz grit and gravel and grey phyllite

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The South Coast of Cornwall

⊡⊡ Fig. 7.9.2.8 The Fal delta, Carrick Roads, showing eroded micro-cliffs at the edge of a salt marsh and a wave-cut ramp in white china clay deposits. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.9 Slope-over-wall coast on the eastern side of Dodman Point. (Courtesy Geostudies.)

pebbles and granules, the sand having come mainly from the Head deposits exposed at the base of the coastal slope. The east-facing coast has a succession of valley-mouth coves at Gorran Haven, Portmellon and Mevagissey, separated by small promontories. Pentewan Beach (> Fig. 7.9.2.11) is wider, formed mainly by sand and brought down by the St. Austell River from a mining hinterland. To the north is Black Head, a rounded promontory, with gorse heath on slopes that decline to cliffs and shore outcrops of massive, hard greenstone. At Charlestown the harbour was built in a valley-mouth cove and the beach of sand and shingle includes pebbles of ballast from the ships that came in to collect china clay. The coast turns eastward past Carlyon Bay (>  Fig.  7.9.2.12),  where a straight and wide beach was ­augmented by the deposition of mining waste (sand and gravel) conveyed by diverting a stream through a tunnel cut in 1842 under the coastal ridge (Everard 1960). This supply has ceased, and the beach is diminishing. Beach accretion in Par Bay accelerated as the result of  the delivery of waste from tin and copper mining in the hinterland by rivers. On the Gribbin Head promontory the coastal slopes descend to cliffs and rocky shores cut in  slates of the Meadfoot Beds, and these extend eastward  past the drowned valley at Fowey to Lantivet Bay. Meadfoot Beds alternate with Dartmouth Slates, the lowest of the Devonian formations seen in Cornwall, as the coast intersects an ­anticlinal structure. The cliffs descend into the deep valley at Polperro, and Talland Bay is at the mouth of two ­converging valleys. Behind Talland Bay the coastal slope on Head deposits descends to low cliffs in

The South Coast of Cornwall

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⊡⊡ Fig. 7.9.2.10 South along Vault Beach, which contains white quartzite pebbles, derived from quartz veins in the Dodman phyllites. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.11 Pentewan Beach, built up by the accretion of sand and gravel waste produced by mining and quarrying in the hinterland, and washed down the St. Austell River. (Courtesy Geostudies.)

the  pink,  purple and grey quartz-veined Dartmouth Slates (> Fig. 7.9.2.13). Looe River flows down a long narrow valley which was invaded by the sea to form a straight inlet, now largely filled with alluvial sediment. To the east the cliffs are mostly cut in dark grey Meadfoot Beds, and remains of a submerged forest are sometimes uncovered at low tide on the Millan­draeth shore. Towards Seaton, steep cliffs of crumbling shales and slates have been cut back quickly by the sea, and there are sectors of arcuate slumping. Between

Seaton and Downderry, cliffs of earthy Head were formerly eroding and slumping because of frequent waterlogging, but the landslides have been stabilised and landscaped. Behind Long Beach there are many small caves angled into the cliffs with sloping roofs on bedding planes (> Fig. 7.9.2.14). Whitsand Bay has a 5 mile gently curving coastline that faces SW, with scrubby coastal slopes descending to low cliffs and protruding ribs of silvery grey Dartmouth Slates (> Fig. 7.9.2.15).

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The South Coast of Cornwall

⊡⊡ Fig. 7.9.2.12 The outlet from the tunnel (arrowed) through which mining waste was delivered to form a beach in Carlyon Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.13 The shore at Talland Bay, showing ridges and furrows in the pink and grey Dartmouth Beds, which strike east-west and dip steeply northward. (Courtesy Geostudies.)

Rame Head projects southward on an outcrop of hard sandstone, and at Penlee Point the coastline swings sharply northward, and becomes sheltered from the strong SW waves. Dartmouth Slates outcrop in the low cliffs and rocky shores NW to Cawsand, on a part of the coast that has forested slopes descending almost to high tide level. These continue round Mount Edgcumbe into the Tamar estuary. This is another drowned valley (ria)

system, branching into St. John’s Lake, the St. Germans estuary, the Tavy and the Plym. It is a “low wave energy coast” because wave action across the estuary is weak, even at high tide, and ineffective as the tide falls. Tidal currents shape intertidal mudbanks, which are typically convex in profile, curving down to the channels that ­persist at low tide. They are backed by segments of salt marsh, generally cliffed along the outer edges. The

The South Coast of Cornwall

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⊡⊡ Fig. 7.9.2.14 Eastward dipping Meadfoot Beds in the cliffs behind Long Beach, Tregantle, showing a cave cut out along bedding planes. (Courtesy Geostudies.)

⊡⊡ Fig. 7.9.2.15 The coast of Whitsand Bay at Sharrow Point, showing ribs of Dartmouth Slate, with east-facing slopes on cleavage planes. (Courtesy Geostudies.)

boundary between Cornwall and Devon follows the Tamar River upstream.

References Arber M (1949) Cliff profiles of Devon and Cornwall. Geogr J 114:191–197

Bird ECF (1998) The coasts of Cornwall. Alexander Publications, Fowey, Cornwall Everard CE (1960) Mining and shoreline evolution near St. Austell, Cornwall. Trans Roy Geol Soc Cornwall 19:199–219 Ranwell DS (1974) The salt marsh to tidal woodland transition. Hydrobiol Bull 8:139–151

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Isles of Scilly

7.10  Isles of Scilly 1.  Introduction The Isles of Scilly, about 48 km SW of Land’s End, are granitic islands produced by partial marine submergence of a granite landscape much like that of the Land’s End Peninsula. The islands consist of a dissected plateau, the higher summits rising between 45 and 49 m (possibly an uplifted marine planation surface of Pliocene age), and there are subdued terraces at lower levels, notably 30–40 m above sea level, representing Pleistocene phases of marine erosion at intermediate levels. About 9,000 years ago, when sea level was 10 m lower than at present, the Isles of Scilly consisted of one large granite island, but the continuing rise of sea level submerged the low-lying parts and left the hills as numerous large and small rocky islands, with associated gravel and boulder beaches, around the inner Sounds. There are about 135 of these islands, including large rocky islets, and Broad Sound is a submerged valley or down-faulted corridor that opens to the SW. Bishop Rock with its prominent lighthouse stands far out to the south–west, beyond numerous irregular rocky skerries (Barrow and Flett 1906). The outer coasts of the islands, especially to the north and west, are exposed to frequent Atlantic storms,

and present bare rugged headlands, islands and stacks. There are many granite tors and buttresses, particularly along the coastal slopes. The granite is divided by joints, mainly vertical or sub-horizontal, but sometimes inclined, and these have influenced the shape of the rocky islands, headlands, cliffs and reefs. Shore platforms are rare, but occur on flat or gently-inclined joint planes. Sandy beaches generally occupy coves and small inlets between granitic headlands. Much of the sand is quartzose, formed by the weathering of granite and swept up from the surrounding shallow sea floor by wave action. Some sand and gravel, including shelly deposits, still arrive on these beaches, but in recent decades beach erosion has become prevalent. Some of the beaches have been shaped into tombolos, linking pairs of islands, such as the sandy barrier on which Hugh Town, on St. Mary’s, is built, linking two former granite islands, and the sandy isthmus on Samson. There are similar intertidal bars of sand and gravel, known as swashways. There are cliff exposures of periglacial Head deposits, which contain rock fragments of local origin in a matrix of sandy clay, and on White Island, St. Martin’s and locally elsewhere there are patches of glacial drift (Mitchell and Orme 1967; Scourse 1991). The glacial

⊡⊡ Fig. 7.10.1 Joint planes in the granite exposed on Inner Head, south–east of Hugh Town. (Courtesy Geostudies.)

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drift on the northern fringes of the Isles of Scilly, ­somewhat south of the general limit of Pleistocene glaciation in Britain, has been attributed to a lobe of Irish Sea ice that spread southward across an emerged sea floor and extended to this area ­during the Anglian (Mindel) Glaciation, about 200,000 years ago.

varying proportions of carbonate derived from shelly organisms. Intervening cliffed sectors have ­slope-over-wall profiles (> Fig. 7.10.3). North–east of St. Mary’s, across Crow Sound, the Eastern Isles are a group of smaller islands arranged in NW–SE groups.

2.  St. Mary’s

3.  St. Martin’s

St. Mary’s, the largest of the Isles of Scilly, consists of a western promontory, Garrison Hill, a former small island attached to the main island by a tombolo (> Fig. 7.10.1). The outlines of the island are strongly influenced by joint patterns in the granite, which also influence the shape of cliffs and tors (> Fig. 7.10.1). The south and east coasts of St. Mary’s show a series of headlands and coves excavated along joint planes, mainly NW–SE. On sheltered parts of the coast, as in Carn Leh Cove and Porth Hellick, the slope in periglacial Head descends almost to high tide level, and there are exposures in low cliffs. In Porth Hellick, the Head deposits are exposed in low cliffs and small caves have been excavated by marine erosion (> Fig. 7.10.2). Beaches on the coast of St. Mary’s are confined to bays and coves, and consist mainly of quartz and felspar derived from weathering and erosion of granite in the bordering cliffs and rocky shores and on the sea floor, together with

St. Martin’s is an elongated island rising to a plateau cut across the granite, probably by marine erosion whenthe sea stood at a higher level. White Island, linked to St. Martin’s by a natural causeway of gravel and boulders (> Fig. 7.10.4), is notable for its glacial drift deposits (till), ­indicating that Irish Sea ice at one stage extended to the northern part of the Isles of Scilly. The glacial drift is prominent in the central part of the island, notably along an incised cleft (>  Fig. 7.10.5). The archipelago of granitic islets, south of St. Martin’s, include boulder beaches, spits and tombolos.

⊡⊡ Fig. 7.10.2 Rounded caves (a metre high, a metre wide and about a metre inset) cut into Head deposits in a cliff bordering Porth Hellick, St. Mary’s, formed by abrasion as granite boulders were agitated by wave action at high tide. (Courtesy Geostudies.)

4.  Tresco Tresco is separated from St. Martin’s by shallow straits. The northern part of the island is a cliff-edged granite ­plateau, but the southern part consists of former small islands attached by sandy depositional features. The

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Isles of Scilly

⊡⊡ Fig. 7.10.3 The slope-over-wall profile on the NE coast of St. Mary’s, with low cliffs cut into the Head deposits that mantle the slope and a rocky shore developed on the underlying granite. (Courtesy Geostudies.)

⊡⊡ Fig. 7.10.4 The bouldery intertidal causeway linking White Island to St. Martin’s. Most of the boulders are of local granite, but there are cobbles and pebbles of flint, basalt and sandstone derived from the island’s glacial drift deposits. (Courtesy Geostudies.)

­ arrow strait between Tresco and Bryher contains relics of n prehistoric field systems and hut circles exposed at low tide. Grassy dunes occupy the southern part of Tresco and cap barriers, which enclose freshwater lakes, on either side of Tresco. West of Tresco, across the strait, is the island of Bryher, which has a sheltered southern coast with rocky shores and bay beaches.

5.  St. Agnes St. Agnes, in the SW of the Isles of Scilly, has rocky northern, western and southern coasts exposed to strong wave action and a more sheltered eastern coast, including the island of Gugh, attached by a sandy tombolo.

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⊡⊡ Fig. 7.10.5 A cliff on White Island exposing glacial drift containing angular fragments of local granite and a variety of erratics including sandstones, flint and volcanic rocks in a yellow-brown earthy matrix. The erratic stones have been derived from regions far to the north and carried here by ice. (Courtesy Geostudies.)

References Barrow G, Flett JS (1906) The Geology of the Isles of Scilly. Memoirs of the Geological Survey, Geological Survey Publication, HMSO, London

Mitchell GF, Orme AR (1967) The Pleistocene deposits of the Isles of Scilly. Quart J Geol Soc London 123:59–92 Scourse JD (1991) Glacial deposits of the Isles of Scilly. In: Ehlers J et al (ed)  Glacial deposits in Great Britain and Ireland. Rotterdam, the Netherlands, pp 291–300

7.11  South Devon The South Devon coast consists of a southern part which resembles the south coast of Cornwall, developed on overthrust Devonian formations that extend from the Tamar estuary across to Torbay, with some headlands of greenstone and older rock formations, notably the Palaeozoic and Pre-Cambrian schists of Start Point and Bolt Tail, which resemble those of the Lizard in Cornwall. The central section is dominated by red cliffs of Triassic sandstone and mudstone within Torbay and from Teignmouth round to Seaton, and the eastern part by the Cretaceous rocks at the head of Lyme Bay. Much of the coast is steep and cliffed. On the southern part there are slope-over-wall profiles similar to those of Cornwall, with vegetated slopes mantled by periglacial Head deposits and basal cliffs cut in hard rock formations, but the receding cliffs on Triassic sandstone are generally vertical (often with outlying stacks), and the coastal slopes in Cretaceous rocks, irregular as the result of landslides. The steep and cliffed coasts are interrupted by incised valleys that have been submerged by the sea to form rias, which have been in varying degree filled with sediment to produce intertidal sandy and muddy

areas, salt marshes and reedswamp, and alluvial valley floors. Wave action is often strong from the SW, and occasionally from the SE and east, particularly in Lyme Bay. Mean spring tide ranges are generally between 4 and 5 m. Beaches are generally sandy, with some gravel in rocky coves, but the shores of Start Bay are bordered by a gently curving beach of brown shingle, predominantly flint pebbles. This includes the barriers that enclose Slapton Ley and smaller lagoons at Beesands.

The South Devon Coast Plymouth Sound is a major ria at the mouth of the River Tamar and its tributaries, notably the Tavy, Lynher and Plym. As the tide falls, wide muddy smooth slopes are exposed, crossed by minor creeks. The muddy sediment has been derived largely from fine-grained material in the Head deposits on bordering slopes, exposed in low cliffs beside the estuary. The entrance to Plymouth Sound is sheltered by a long breakwater.

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⊡⊡ Fig. 7.10.5 A cliff on White Island exposing glacial drift containing angular fragments of local granite and a variety of erratics including sandstones, flint and volcanic rocks in a yellow-brown earthy matrix. The erratic stones have been derived from regions far to the north and carried here by ice. (Courtesy Geostudies.)

References Barrow G, Flett JS (1906) The Geology of the Isles of Scilly. Memoirs of the Geological Survey, Geological Survey Publication, HMSO, London

Mitchell GF, Orme AR (1967) The Pleistocene deposits of the Isles of Scilly. Quart J Geol Soc London 123:59–92 Scourse JD (1991) Glacial deposits of the Isles of Scilly. In: Ehlers J et al (ed)  Glacial deposits in Great Britain and Ireland. Rotterdam, the Netherlands, pp 291–300

7.11  South Devon The South Devon coast consists of a southern part which resembles the south coast of Cornwall, developed on overthrust Devonian formations that extend from the Tamar estuary across to Torbay, with some headlands of greenstone and older rock formations, notably the Palaeozoic and Pre-Cambrian schists of Start Point and Bolt Tail, which resemble those of the Lizard in Cornwall. The central section is dominated by red cliffs of Triassic sandstone and mudstone within Torbay and from Teignmouth round to Seaton, and the eastern part by the Cretaceous rocks at the head of Lyme Bay. Much of the coast is steep and cliffed. On the southern part there are slope-over-wall profiles similar to those of Cornwall, with vegetated slopes mantled by periglacial Head deposits and basal cliffs cut in hard rock formations, but the receding cliffs on Triassic sandstone are generally vertical (often with outlying stacks), and the coastal slopes in Cretaceous rocks, irregular as the result of landslides. The steep and cliffed coasts are interrupted by incised valleys that have been submerged by the sea to form rias, which have been in varying degree filled with sediment to produce intertidal sandy and muddy

areas, salt marshes and reedswamp, and alluvial valley floors. Wave action is often strong from the SW, and occasionally from the SE and east, particularly in Lyme Bay. Mean spring tide ranges are generally between 4 and 5 m. Beaches are generally sandy, with some gravel in rocky coves, but the shores of Start Bay are bordered by a gently curving beach of brown shingle, predominantly flint pebbles. This includes the barriers that enclose Slapton Ley and smaller lagoons at Beesands.

The South Devon Coast Plymouth Sound is a major ria at the mouth of the River Tamar and its tributaries, notably the Tavy, Lynher and Plym. As the tide falls, wide muddy smooth slopes are exposed, crossed by minor creeks. The muddy sediment has been derived largely from fine-grained material in the Head deposits on bordering slopes, exposed in low cliffs beside the estuary. The entrance to Plymouth Sound is sheltered by a long breakwater.

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South Devon

The coast at Plymouth is on Devonian limestone, which also outcrops on Mount Batten Point. From this Point coastal slopes rise to Staddon Point, where there are cliffs and basal rock ledges cut in the Lower Devonian Staddon Grits. To the south the underlying Meadfoot Beds emerge in slaty cliffs and irregular shore platforms in Bovisand Bay, and the Dartmouth Slates outcrop in cliffs and a wide shore platform extending round to Wembury Point. The slope-over-wall coastline on Dartmouth Slates then trends eastward into the Yealm estuary below Wembury. The Yealm estuary is another ria, joined by a tributary at Newton Ferrers. On the southern side, the Headmantled coastal slopes descend almost to high tide level, but basal cliffs in Dartmouth Slates develop past Mouthstone Point and increase to Gara Point. A series of coves and rocky promontories continues eastward to the mouth of the Erme estuary. The Erme estuary (> Fig. 7.11.1) has much inwashed sand. Only a small meandering river channel remains between sandflats at low tide, and upstream the sand gives place to muddy sediment. Beacon Point marks the beginning of Bigbury Bay, where slope-over-wall profiles border little promontories and sandy valley-mouth coves, as at Westcombe and Challaborough. Bigbury-on-Sea is linked to Burgh Island, a small island with Head-mantled slopes and rocky shores, by an intertidal sand bar. The cliffs are cut in Dartmouth Slates. Sand has been washed onshore from Bigbury Bay,

forming dunes at Bantham. There are offset spits at the mouth of the River Avon. A low slope-over-wall coast extends along the eastern side of Bigbury Bay, where the Dartmouth Slates disappear beneath Meadfoot Beds, and at Thurlestone there is a coastal outcrop of Triassic sandstone running out to Warren Point. Rocky ledges on Meadfoot Beds extend along the cliffs of Great Ledge to the two coves, Outer Hope and Inner Hope. The coast steepens from Hope Cove to Bolt Tail, the beg­­inning of a high section of steep coast on Palaeozo­ icand Pre-Cambrian mica schists that runs SE to Bolt Head, with bolder promontories on harder hornblende schist. There are tors on the coastal slope above the cove at Fernyhole Point. The Warren is a broad scrubby coastal slope with slumping in the Head deposits, and at Bolt Head the coast swings northward and steep, high cliffs extend past Sharp Tor and successive valley-mouth coves opening to sandy beaches, as at South Sands. Salcombe is built on a slope bordering a deep ria, extending inland as the Kingsbridge estuary. To the east the bold coast resumes, with steep, often irregular Head slopes and basal cliffs in hornblende schists out to Prawle Point. The coast then faces SE, and consists of Headaproned terraces backed by bluffs marking degraded Pleistocene cliffs and fronted by low cliffs and wide rocky shore platforms. Periglacial drift (Head) is exposed in bluffs behind Lannacombe Bay and in conical stacks.

⊡⊡ Fig. 7.11.1 The sanded estuary of the River Erme at low tide. (Courtesy Geostudies.)

South Devon

An incised valley opens on to Lannacombe Beach and to the east, the coastal terraces fade into a steepening slope, undercut by cliffs along to Start Point, a craggy headland of mica schist, with slopes with periglacial Head descending to cliffs which are higher on the more exposed southern side than on the northern flank facing Start Bay (> Fig. 7.11.2). The steeply sloping coast curves north from Start Point past the ruined village of Hallsands (> Fig. 7.11.3). In the nineteenth century the village was protected by a high, wide shingle beach, but cliff recession has been rapid here following the dredging of shingle between 1897 and 1902 to obtain gravel for the construction of Plymouth dockyard. The Start Bay beaches are relict, receiving no shingle under present conditions, so that the losses have not been compensated, and erosion is continuing. The floor of Start Bay is sandy, but there is shingle out  on Skerries Bank, which may be a submerged barrier  beach. On the coast the shingle beach is interrupted  by  Tinsey Head, but resumes to enclose a small lagoon at Beesands. To the north is a quarried headland of  greenstone, Limpet Rocks, with the village of Tor­­ cross beyond. A barrier beach up to 140 m wide, consisting of  fine flint shingle encloses a lagoon, Slapton Ley (> Fig. 7.11.4), which has a controlled outlet at the southern end. The Ley is almost fresh, receiving salt water and spray only during easterly storms, when waves may wash over the barrier beach and sweep shingle across into the lagoon. The barrier beach is thus migrating landward, narrowing the lagoon. There is no local source of flint ⊡⊡ Fig. 7.11.2 Coastal slopes on Head deposits, looking towards Start Point. (Courtesy Geostudies.)

7.11

pebbles, which may have been swept shoreward from the  sea floor  during  the Late Quaternary marine transgression. The northern part of Slapton Ley is occupied by reedswamp. At Pilchard Cove the Meadfoot Beds come to an end as the underlying Dartmouth Beds rise. There are rocky ribs and coves, the largest of which contains Blackpool Sands, and occasional headlands of greenstone, as on Matthews Point, on to the mouth of the River Dart. The long estuary of the River Dart is a ria cut into Dartmouth Slates, and is tidal to above Totnes. As it opens to the sea the northern side of the ria has steep wooded slopes extending seaward past Mill Bay Cove and New­ foundland Cove to headlands on igneous rock at Inner and Outer Froward Points. A steep coast with slope-­ over-wall profiles then continues northward with many small headlands and coves such as Scabbacombe Sands (> Fig. 7.11.5). Man Sands is a lagoon enclosed by a ­scrubcovered barrier that was breached during a flood in 2008. The coast then runs out to a small Devonian limestone promontory at Sharkham Point, and St Mary’s Bay is cut in slates. Berry Head rises to a plateau 60 m above sea level and its northern cliff has been cut back by quarrying. The coast runs westward along the northern side of the Brixham peninsula, with wooded bluffs on the southern coast of Tor Bay. Tor Bay has been excavated in soft Permo-Triassic sandstone between harder Devonian rocks to the south and north. Elbery Cove contains a beach of well-rounded cobbles, and Broad Sands are bordered by cliffs cut in grey

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South Devon

⊡⊡ Fig. 7.11.3 The ruined village at Hallsands. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.4 Slapton Ley and the barrier beach north of Torcross. (Courtesy Geostudies.)

Devonian limestone, passing to red Triassic sandstones on  the northern side. Roundham Head is fringed by cliffs of red sandstone which continue past Paignton and Hollicombe to Torquay. The Torquay peninsula runs out to Hope’s Nose, where the coastal slopes descend to low cliffs and shore platforms in Devonian limestone, with shingle beaches in coves, passing to greenstone at Black Head and Long Quarry Point. On the northern side of the peninsula Permian sandstones and breccias begin below Babbacombe, where

they are faulted against pale weathered Upper Devonian sandstones. Shingle beaches are derived from the breccias, and the shore is littered with fallen blocks of sandstone north of Oddicombe Beach. The Devonian limestone returns at Petit Tor Point and Watcombe Head has red sandstone cliffs. At the Ness south of Teignmouth, rounded Permian sandstone blocks have been washed out of a sandy matrix in the cliff. At Teignmouth the Teign estuary has wide muddy sandflats and some gravelly shoals at low tide, patches of salt

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⊡⊡ Fig. 7.11.5 Scabbacombe Sands occupy a cove on the northern side of Scabbacombe Head, near Dartmouth. (Courtesy Geostudies.)

marsh, minor beaches and forelands, and reed swamp upstream near Newton Abbot. A spit has grown south almost across the mouth of the Teign estuary. North of Teignmouth red sandstone and conglomerate cliffs extend to Dawlish and Dawlish Warren - the coast having been modified by the construction of the Great Western Railway. Dawlish Warren is an unusual double spit, consisting of the Inner and Outer Warrens with sand dunes and Greenland Lake occupying the intervening swale (Kidson 1950). There are shoals of sand off Exmouth and sand and gravel in the Exe estuary to the north. East of Exmouth cliffs of red sandstone begin at Orcombe and extend past Sandy Bay to Littleham Cove. This marks the start of a beach of ovoid cobbles that extends to Budleigh Salterton, and has been derived from cobble beds that descend through the cliff (> Fig. 7.11.6). East of Budleigh Salterton, the River Otter estuary was once a long narrow marine inlet, but is now occupied by salt marsh with much Spartina, backed by rushes and reeds. Bluffs (degraded cliffs) run up the eastern side of the Otter valley, and the shallow river mouth is constricted by a recurved shingle spit. Cliffs with shore platforms on the overlying red Triassic sandstone start at Otterton Point, where Otterton Ledge is a protruding sandstone bench. At Ladram Bay, thick almost flat bedded red sandstones form vertical cliffs 20–25 m high, and have been dissected along joint planes to form stacks, surrounded by a sandy beach with ­scattered

pebbles and coves with pocket beaches of shingle. The sandstone cliffs increase in height rapidly to High Peak (157 m) and Peak Hill (156 m), which are truncated by steep cliffs with angular vertical grooves and spurs on the cliff face (> Fig. 7.11.7). High cliffs of red Triassic marl and sandstone with buttresses and a basal apron of slumped sediment run behind the beach to the promontory at Sidmouth. The shingle beach here had become much depleted by 1995, when the shore was protected by boulders overlain by gravel brought from Woodbury quarry to form a ­nourished beach, which was held between three groynes and sheltered by two en echelon offshore breakwaters at the western end, built to break up SW storm waves. The River Sid percolates through a shingle bank to the sea. To the east are cliffs in red marl, sandstone and breccia, dipping gently eastward and fronted by shingle beaches interrupted by salients of sandstone and boulders known locally as ebbs. Rock falls occur frequently. Successive headlands have a capping of Upper Greensand, and are separated by deep valleys, as at Weston Mouth and Branscombe. Shingle beaches, consisting mainly of flint from the Chalk and chert from the Upper Greensand, occupy compartments along the coast, and typically show lateral gradation, with larger pebbles and cobbles towards the eastern ends (Bird 1989). To the east the cliff rises, the slope steepening as it passes from Triassic mudstone to the overlying Upper

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⊡⊡ Fig. 7.11.6 Gravel beds in the cliff west of Budleigh Salterton are the source of distinctive ovoid beach cobbles. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.7 The gullied cliff at Peak Hill, Sidmouth. (Courtesy Geostudies.)

Greensand and Chalk. The shingle beach ends at a landslide under Hooken, forming a hummocky undercliff with buttresses and blocks that have fallen from the upper cliff of Chalk. This descends to the coast at Beer Head, where the cliffs turn back into the shingle cove at Beer (> Fig. 7.11.8). In Seaton Bay a shingle beach fronts cliffs of red and grey marl, interrupted by a wooded chine, and then the esplanade begins. Seaton has a wide shingle beach, and the

sea wall forming the esplanade has tidal gates built in 1980 after a 1979 storm surge flooded the seafront and swept shingle over the road. At the eastern end of Seaton esplanade the River Axe reaches the sea, having been diverted eastward by a longshore shingle spit. The Axe estuary is broad at  high tide, and fringed by salt marshes backed by reedswamp. Triassic red marls form the cliffs from Axmouth to  Haver Cliff, the pink rock declining beneath white

South Devon

l­ imestones, then Blue Lias and Black Ven Marls. Cretaceous formations (Chalk, Upper Greensand and Gault) ­outcrop above, behind a subsided fringe. This is the start of the great Axmouth-Dowlands Landslide (Landslip), which took place at Christmas 1839 (Arber 1940). Masses of permeable Chalk and Upper Greensand slid seaward over Gault clay when the interface was lubricated and slurried by percolating groundwater. Behind the subsided area, an upper cliff of Chalk forms an abrupt edge to the coastal plateau. Luxuriant scrub ⊡⊡ Fig. 7.11.8 Chalk cliffs at Beer, looking towards Beer Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.11.9 The cliffs at Pinhay Bay, looking west towards the Landslip at Humble Point. (Courtesy Geostudies.)

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and woodland vegetation occupies a subsided area of dislocated blocks and chasms, with numerous fractures that are active, especially after wet weather, indicating that subsidence is continuing. The shore in front of the Axmouth-Dowlands Landslide is strewn with boulders, has irregular outcrops of tilted strata and segments of gravelly beach as in Charton Bay. Tumbled rocks extend along the shore past Humble Point to Pinhay Bay, where there is a basal cliff in Rhaetic White Lias (> Fig.  7.11.9). Steeply-sloping cliffs

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run on towards Seven Rock Point, where the basal cliff is in the underlying Blue Lias. The cliffs decline into scrubby bluffs at Devonshire Head which formerly protruded seaward and the shore  platform shows limestone layers forming scarplets between shale bands. The Devon-Dorset boundary crosses the coast half a mile west of The Cobb at Lyme Regis.

References Arber MA (1940) The coastal landslips of south-east Devon. Proc Geol Assoc 51:257–271 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Kidson C (1950) Dawlish Warren: a study of the evolution of sand spits across the mouth of the River Exe in Devon. Trans Inst Br Geogr 16:67–80

7.12  Dorset 1.  Introduction The Dorset coast is dominated by Jurassic formations, overlain by Cretaceous rocks including the Chalk, which forms cliffs between White Nothe and Lulworth Cove, in Worbarrow Bay, and on Ballard Down. The Chalk outcrop narrows eastward as the dip increases to form the Purbeck Ridge, which runs along the northern fringe of the Purbeck Anticline. To the north Lower Tertiary (Eocene) formations occupy the lowland around Poole Harbour and along the coast of Bournemouth Bay (Bird 1995). The present coastline of Dorset is largely the outcome of erosion and deposition during the past 6,000 years, since the sea reached its present level as a sequel to the worldwide Late Quaternary marine transgression. There are only a few surviving Pleistocene features, such as the Portland Raised Beach. It is likely that, as the Holocene sea rose to its present level, high coastal sectors presented steep, formerly periglaciated slopes, but these were soon undercut by marine erosion and converted into receding cliffs, the crests of which undulate across the truncated hills and valleys. In general, promontories occur where harder rocks outcrop at sea level, while coves and bays have been excavated in softer outcrops, particularly where these have already been lowered by stream erosion. In detail, lines of weakness such as joints and faults have been exploited to form clefts and caves, and locally to isolate stacks. Cliff profiles are related to varying rock resistance and permeability, so that the Chalk, Greensands and Jurassic limestones and sandstones all form bold outcrops, especially where they have been hardened as the result of intensive folding. However, in addition to geological factors it is necessary to take account of the effects of weathering and rainwash, as well as marine processes and con­­trasts between sheltered sectors and those exposed to high wave energy.

Cliff recession has been accompanied by landslides of various kinds, where masses of rock, weathered material and soil have slipped from high to low ground (Arber 1973; Brunsden and Jones 1972). They have occurred ­particularly where a capping of Chalk and Upper Green­ sand has collapsed over underlying outcrops of the less permeable Gault and Lias clays, lubricated to mudslides by seeping groundwater, as on Ware Cliffs and White Nothe. In a similar way the Portland Beds have slid down over Kim­meridge Clay, as at St. Alban’s Head, Hounstout Cliff, Gad Cliff and on the northern slopes of the Isle of Portland. The movements have been over a clearly definable surface, either translational (down a plane) or rotational (over a concave curve). On soft fine-grained sediment, mudslides usually lobate and elongated, move slowly downslope, torn by shearing and tension cracks. As they subside they leave arcuate cliff-backed hollows. More rapid movements, especially during very wet weather, include mudflows (where unconsolidated materials soaked with water move as a viscous mass), topples (where lumps of rock and soil break away and subside) and rock falls (when detached material drops from a cliff and forms a scree of talus). Shore platforms have been cut by marine erosion as cliffs recede. Some are seaward-sloping intertidal abrasion platforms while others are subhorizontal structural ledges on the surface of resistant strata outcropping at the cliff base. West of Portland Bill the Dorset coast is exposed to ocean swell from the Atlantic Ocean and strong wave action generated by the prevailing south-westerly winds; to the east it is at first sheltered by the Isle of Portland, but exposure increases again beyond White Nothe. There is occasionally strong wave action from the south and southeast, most evident on east-facing shores in Weymouth Bay, and between Durlston Head and Poole. Mean spring tide ranges are smaller on the Dorset coast than in south-west or south-east England. At Lyme Regis

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run on towards Seven Rock Point, where the basal cliff is in the underlying Blue Lias. The cliffs decline into scrubby bluffs at Devonshire Head which formerly protruded seaward and the shore  platform shows limestone layers forming scarplets between shale bands. The Devon-Dorset boundary crosses the coast half a mile west of The Cobb at Lyme Regis.

References Arber MA (1940) The coastal landslips of south-east Devon. Proc Geol Assoc 51:257–271 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Kidson C (1950) Dawlish Warren: a study of the evolution of sand spits across the mouth of the River Exe in Devon. Trans Inst Br Geogr 16:67–80

7.12  Dorset 1.  Introduction The Dorset coast is dominated by Jurassic formations, overlain by Cretaceous rocks including the Chalk, which forms cliffs between White Nothe and Lulworth Cove, in Worbarrow Bay, and on Ballard Down. The Chalk outcrop narrows eastward as the dip increases to form the Purbeck Ridge, which runs along the northern fringe of the Purbeck Anticline. To the north Lower Tertiary (Eocene) formations occupy the lowland around Poole Harbour and along the coast of Bournemouth Bay (Bird 1995). The present coastline of Dorset is largely the outcome of erosion and deposition during the past 6,000 years, since the sea reached its present level as a sequel to the worldwide Late Quaternary marine transgression. There are only a few surviving Pleistocene features, such as the Portland Raised Beach. It is likely that, as the Holocene sea rose to its present level, high coastal sectors presented steep, formerly periglaciated slopes, but these were soon undercut by marine erosion and converted into receding cliffs, the crests of which undulate across the truncated hills and valleys. In general, promontories occur where harder rocks outcrop at sea level, while coves and bays have been excavated in softer outcrops, particularly where these have already been lowered by stream erosion. In detail, lines of weakness such as joints and faults have been exploited to form clefts and caves, and locally to isolate stacks. Cliff profiles are related to varying rock resistance and permeability, so that the Chalk, Greensands and Jurassic limestones and sandstones all form bold outcrops, especially where they have been hardened as the result of intensive folding. However, in addition to geological factors it is necessary to take account of the effects of weathering and rainwash, as well as marine processes and con­­trasts between sheltered sectors and those exposed to high wave energy.

Cliff recession has been accompanied by landslides of various kinds, where masses of rock, weathered material and soil have slipped from high to low ground (Arber 1973; Brunsden and Jones 1972). They have occurred ­particularly where a capping of Chalk and Upper Green­ sand has collapsed over underlying outcrops of the less permeable Gault and Lias clays, lubricated to mudslides by seeping groundwater, as on Ware Cliffs and White Nothe. In a similar way the Portland Beds have slid down over Kim­meridge Clay, as at St. Alban’s Head, Hounstout Cliff, Gad Cliff and on the northern slopes of the Isle of Portland. The movements have been over a clearly definable surface, either translational (down a plane) or rotational (over a concave curve). On soft fine-grained sediment, mudslides usually lobate and elongated, move slowly downslope, torn by shearing and tension cracks. As they subside they leave arcuate cliff-backed hollows. More rapid movements, especially during very wet weather, include mudflows (where unconsolidated materials soaked with water move as a viscous mass), topples (where lumps of rock and soil break away and subside) and rock falls (when detached material drops from a cliff and forms a scree of talus). Shore platforms have been cut by marine erosion as cliffs recede. Some are seaward-sloping intertidal abrasion platforms while others are subhorizontal structural ledges on the surface of resistant strata outcropping at the cliff base. West of Portland Bill the Dorset coast is exposed to ocean swell from the Atlantic Ocean and strong wave action generated by the prevailing south-westerly winds; to the east it is at first sheltered by the Isle of Portland, but exposure increases again beyond White Nothe. There is occasionally strong wave action from the south and southeast, most evident on east-facing shores in Weymouth Bay, and between Durlston Head and Poole. Mean spring tide ranges are smaller on the Dorset coast than in south-west or south-east England. At Lyme Regis

Dorset

the range is 3.7 m, and at Bridport 3.5 m. The tides diminish east of Portland Bill to 1.9 m at Portland, 2.0 m in Lulworth Cove and 1.7 m between Swanage and Bournemouth. There are the complications of prolonged low tide in Portland Bay and prolonged high tide in Bournemouth Bay - the latter facilitating wave attack on cliffs. The effects of marine erosion and deposition have been to simplify coastal outlines, so that only the most resistant rocks persist as promontories, such as the massive limestones of Portland Bill and the southern shores of Purbeck. Lyme Bay has evolved from a series of small bays with irregular headlands to the relatively smooth coastline that now runs from Torquay to Portland Bill as the result of cliff recession and the deposition of Chesil Beach. The coastline has become adjusted to fit the patterns of waves that move in across Lyme Bay. Between Portland Bill and the Isle of Purbeck there are still minor rocky promontories between small curving bays with beaches, but in Bournemouth Bay the cliffed coastline cut in softer Eocene rocks has become more open and smoothly curved, with some sandy deposition on the western shores. As these cliffs receded, small streams entrenched steep-sided valleys known as chines (Bury 1920). The beaches of Dorset are generally of local derivation, consisting of sand and gravel produced by erosion of nearby cliffs and shore outcrops, and distributed along the coast by wave action. But, as is evident on Chesil Beach, some beach material has been washed up from the sea floor. Most of the beaches are dominated by flint and chert shingle, with locally derived limestone pebbles and sandy material from weathered sandstones, especially in Bou­ rnemouth Bay, where the sandy beaches have been derived from Eocene formations. Some beaches are still receiving sand and gravel eroded from nearby cliffs or the sea floor but others, like Chesil Beach now have little if any sediment supply. Very little beach material has come from rivers, which have been mainly delivering fine-grained sediment (silt and clay) to the coast. In general the predominance of wind and wave action from the south-west has caused beach material to drift eastward, as shown by the accretion of beaches on the western sides of natural obstacles such as landslide lobes under Golden Cap and man-made ob­­ stacles such as breakwaters and groynes, like the Cobb at Lyme Regis. Estuarine inlets have been formed by Late Quaternary marine submergence of valley mouths, notably in Poole and Christchurch Harbours, which have been modified by Holocene deposition to alluvial plains and areas of salt marsh, mudflats and shoals.

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2.  The Dorset Coastline East of Devonshire Head the cliffs have been cut in the layered strata of the Blue Lias and the intertidal area consists of broad structural ledges of limestone where the intervening blue clay horizons have been washed out (>Fig.  7.12.1). Chippel Bay occupies a gentle syncline, with ledges rising again towards the Cobb, the stone breakwater that shelters the ancient harbour of Lyme Regis. Early in the nineteenth century the coast at Lyme Regis had a wide curving beach of brown shingle in front of scrubby slumping bluffs (Wanklyn 1922). A stone and concrete promenade was built at Cobb Gate, and later extended to protect the coastal slopes from wave erosion, and there has been recurrent slumping in the coastal slopes of Middle Lias. Attempts are being made to stabilise the coastal slopes in Lyme by draining off groundwater. Eastward drifting of beach material has been prevented by the Cobb, and the shingle beach of the Lyme seafront has been gradually depleted. It was restored artificially in 2007. At the eastern end of Lyme seafront the coastline curves behind Church Cliffs towards Black Ven. The Blue Lias is overlain by marls, and there is an upper, receding cliff in Gault and Upper Greensand. There have been frequent landslides, the Lias marls and capping Gault having subsided irregularly in front of the receding cliff of Upper Greensand (>Fig. 7.12.2). Cracks develop at the top of the cliff, which recede as blocks calve away and collapse (>Fig. 7.12.3). The cliffs descend towards Charmouth where the mouth of the River Char is encumbered by shingle, and there is a looped intertidal shingle bar off the river mouth. The cliffs ascend steeply towards Stonebarrow Hill, which has a high arcuate landslide back-scar (a curved slipplane) in Upper Greensand, overlooking an amphitheatre of slumped material, Cains Folly, in which rotational slides are moving down towards a 40–60 m lower cliff, where mudslides spill over and down to the shore (>Fig.  7.12.4). This receding coast may be in ‘dynamic equilibrium’ over periods of a century or so, meaning that there have been similar associations of upper and lower cliff and intervening slumped topography, maintaining the transverse profile as the coast retreated (Brunsden and Jones 1976, 1980). Golden Cap (191 m) is the highest coastal cliff in southern England. Its summit is one of many flat-topped ridges, dissected remnants of the West Dorset plateau, and the cliff has yellow-brown Upper Cretaceous rocks over the grey Liassic clays and limestones. There is landslide topography, with slumping slopes in Lias beneath the Upper Greensand capping. Mud lobes have descended to the shore, separating pocket beaches of shingle in small

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⊡⊡ Fig. 7.12.1 Limestone ledges and beach boulders on the shore at Devonshire Head, looking east towards the Cobb at Lyme Regis. The dark marl horizon can be seen below a limestone layer in the scarplet. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.2 The slumping cliff at Black Ven, showing the 2008 landslide undercut at the base. (Courtesy Geostudies.)

coves, and the slumped material has been cut back to leave boulder festoons on the shore. At Seatown the cliffs decline to the mouth of the River Winniford, which percolates out through stormpiled shingle on a beach that is low, flat and sandy with only sparse fine shingle to the west but a higher, steeper and coarser shingle beach to the east. This lateral

g­ radation is characteristic of the beaches of Lyme Bay (Bird 1989). Pebbles have been quarried from this beach for use in ornamental walls and pavements (>Fig. 7.12.5). The cliff east of Seatown is of slumping blue clay, and as it rises towards Ridge Hill the slope steepens on overlying sandstones. Doghouse Hill (132 m) has a small ­capping

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⊡⊡ Fig. 7.12.3 Capping of weathered (oxidised) Upper Greensand on the receding upper cliff at Black Ven. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.4 A basal fan of slumped debris that spilled over the lower cliff east of Charmouth during the wet spring of 2001. It was subsequently cut back behind an abrasion ramp cut in shale. (Courtesy Geostudies.)

of yellow Bridport Sands. Each of the hill crests carries cherty gravel in a brown earthy matrix, and the cliff crest shows that the slopes are mantled by similar material washed downhill. A coombe separates Doghouse Hill from the higher Thorncombe Beacon (157 m), and the little valley at Eype opens between low slumping clay cliffs. Recurrent slumping on West Cliff has been countered by the building of a promenade. The West Bay harbour breakwaters were built as piers in 1823–1825 and extended as a marina in 2003.

The vertical East Cliff and Burton Cliff, on either side of the Bride valley below Burton Bradstock, are cut in Bridport Sands, here almost horizontal calcareous sandstone layers with intervening softer yellow sand. Where there have been recent cliff falls there are basal talus cones, and the cliff face is smooth, but intervening areas show protruding structural ledges of calcareous sandstone etched out by weathering (>Fig. 7.12.6). At Burton Beach the cliffs end abruptly, but Chesil Beach continues as a shingle bank about 150 m wide,

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⊡⊡ Fig. 7.12.5 Seatown, looking west along a shingle beach towards Golden Cap. (Courtesy Geostudies.)

fronting marshy meres, and later the lagoon known as The Fleet (>Fig. 7.12.7). Its crest gradually increases in height from about 6 m above low tide level at Burton Cliff to 14 m at Chiswell, at the SE end. Chesil Beach is a vast accumulation mainly of flint and chert shingle, with some Triassic quartzites, a few pebbles of granite, quartz and metamorphic rock, all derived from formations that outcrop in South Devon. At the south-eastern end there is an increasing proportion of limestone pebbles and cobbles derived from the cliffs of the Isle of Portland. It is essentially a relict formation (no longer receiving much beach material) and probably originated as an offshore barrier which was driven landward during the Late Quaternary marine transgression in such a way that the inner shores of The Fleet have never been exposed to erosion by the waves of the open sea. There is lateral grading of the beach material, the longest diametre of pebbles at West Bay being about 1 cm, increasing gradually to 2–3 cm at Abbotsbury and 10–15 cm at Portland. This is due largely to sorting of the

a­ vailable shingle by size, for attrition of flint and chert pebbles is very slow. Long­shore drifting occurs when waves arrive obliquely to the beach, and it may be that the sorting is the outcome of the contrast between stronger westerly and SW waves, which can move all grades of pebbles, and the weaker SE waves, which take back only the smaller material. This has proceeded over a long period, during the landward migration of Chesil Beach in Holocene times (Carr and Blackley 1974). The Fleet is a shallow, brackish tidal lagoon, separated from the sea by Chesil Bank and having a tidal entrance from Portland Bay at the SE end. It is about 8 miles (13 km) long, from Abbotsbury Swannery to the tidal entrance at Small Mouth, and varies in width from about 75 m at The Narrows to nearly a kilometre in Butterstreet Cove. Tide range increases south-eastward to Small Mouth, where there is prolonged ebb because sea seeps in through Chesil Beach while the tide is still falling in Portland Harbour, especially when sea level in Lyme Bay is raised by westerly winds. At low tide extensive shoals of muddy sand, bearing seagrasses, are exposed in the southeast, while the north-west, towards Abbotsbury Swannery, is less tidal and somewhat freshened by inflowing streams, so that reedswamp has invaded the shallows. Chesil Beach comes to an end below West Weare, and vertical cliffs in Portland Limestone extend south towards Portland Bill (>Fig. 7.12.8). At Portland Bill storm-washed structural ledges of Portland Stone end in cliffs that descend directly into deep water, and Pulpit Rock is a joint-bounded stack rising from a structural ledge. The Raised Beach, much disturbed by quarrying, is banked on a ledge of Portland Limestone, and forms a gently sloping terrace that rises from 7 to 18 m above present sea level and extends for about a mile to the NE. It was formed when sea level was about 12 m higher than it is now, and is thought to comprise two emerged beaches, formed 210,000 years and 125,000 years ago (Davies and Keen 1985). Offshore is the Tide Race, where the ebb and flow of tides interacts with waves to produce rough water over the shingle shoal known as The Shambles. The east coast of the Isle of Portland is more sheltered from strong wave action than the west coast, but the steep sharp-crested coastal slope of Portland Stone increases in altitude northward to East Weare. Portland Bay has been enclosed by breakwaters to form Portland Harbour, and piers have been built on either side of the harbour at the mouth of the River Wey. Weymouth promenade, to the east, is fronted by a wide curving sandy beach, the upper part of which becomes gravelly near the clock tower. The shingle beach then becomes wider and higher, usually with several berms,

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⊡⊡ Fig. 7.12.6 East Cliff, near Bridport, is cut in horizontally bedded sandstones. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.7 The crest of Chesil Beach near Abbotsbury, with the Fleet lagoon on the left. (Courtesy Geostudies.)

continuing in front of the concrete sea wall that borders the Lodmoor lowland. The beach has been augmented by dumping shingle to protect the sea wall. There are cliffs in Oxford Clay to the east, and Redcliff Point is bordered by mudflows, which descend calving, crumbling, flaking and slurrying to form lobate tongues that advance over the shingle beach on the eastern side. This beach consists of brown flint and chert shingle,

g­ enerally coarser to the west and finer to the east, in contrast with Chesil Beach and the Lyme Bay beaches. This is probably because this part of the coast is sheltered from the prevailing westerly and SW waves by the Isle of Portland, so that here the SE waves are dominant, and have moved the larger pebbles westward. Kimmeridge Clay outcrops in the upper part of the sloping cliff at Black Head, and a slumping cliff declines to

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⊡⊡ Fig. 7.12.8 Cliffs in Portland Limestone on the coast of Portland Bill. (Courtesy Geostudies.)

the valley at Osmington Mills. Cliffs continue, subsiding behind the shingle beach at Ringstead, and then rising again to Burning Cliff (so-called because a fire ignited spontaneously in the Kimmeridge oil shale in 1826, and persisted for several years). This overlooks a widening area of scrubby landslide topography on the Kimmeridge Clay, which forms low cliffs and concave ramps immediately behind the beach. Below Holworth House the Kimmeridge Clay is overlain by a disintegrating cliff of Portland Sand, capped by Portland Stone then Purbeck Beds, all dipping north-west. These Upper Jurassic formations are overlain unconformably by the more gently south-eastward dipping Gault, Upper Greensand and Chalk, which descend through the upper cliffs to White Nothe. The cliffs to the east are cut in frost-shattered Chalk, rising and falling as they intersect ridges and coombes. Beyond Swyre Head, is the irregular promontory of Durdle Door. At the eastern end of St. Oswald’s Bay the Chalk passes inland behind a valley cut in the underlying softer formations. This climbs behind the Purbeck and Portland ridge on Dungy Head to meet another valley that descends into Lulworth Cove. This occurs where converging valleys cut into the soft Wealden sands and clays have been scoured out by marine erosion behind a coastal ridge of steeply dipping Purbeck and Portland Limestone. It is backed by a steepened slope cut into the Chalk escarpment after the outer rampart of Purbeck and Portland Limestone had been breached by the sea (>Fig. 7.12.9). Along the coast to the east, a steep escarpment of Portland Stone is capped by Purbeck Beds. The outer ­rampart of Portland Stone breaks up into Mupe Rocks,

and in Mupe Bay the cliffs curving behind a beach of sand and gravel show a sequence from the Purbeck Beds through the crumbling Wealden to the Lower Greensand, partly obscured by Chalk talus, the Gault and Upper Greensand and then a sudden rise to the high Chalk cliff cut into the escarpment. The cliff crest descends just as abruptly to Arish Mell, where cliff recession has intersected the mouth of a steep-sided but gently declining dry valley to form a gap. Flower’s Barrow is an Iron Age hilltop camp truncated by cliff recession, and the cliff at Cow Corner rises sharply, with ribs of chalk jutting out between clefts and caves. Worbarrow Bay is an enlarged and more open version of Lulworth Cove, exposing the Wealden Beds cut back through the Greensands and Gault to the Chalk. At the eastern end of the bay is Worbarrow Tout, a steep promontory much like Durdle Door. The Chalk ridge runs inland as the Purbeck Hills, and from Pondfield Cove the Portland and Purbeck Beds ascend rapidly to outcrop high in Gad Cliff. At the eastern end is Tyneham Cap, where the Portland Stone becomes a well-defined escarpment that runs behind a broad amphitheatre of Kimmeridge Clay, and round to Swyre Head. Behind Brandy Bay are cliffs cut in the layered limestones and blue shales of the Kimmeridge Clay, fronted by structural ledges developed on resistant limestone layers, as at Broad Bench (>Fig. 7.12.10). Undulating cliffs to the east are cut in the Kimmeridge Clay dipping eastward. Successive limestone bands des­ cend to the cliff base, forming slight salients, running out  across the shore as ledges, some of which have

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⊡⊡ Fig. 7.12.9 Lulworth Cove. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.10 Broad Bench, a structural ledge. (Courtesy Geostudies.)

­ isintegrated into boulder beaches. The deeply incised d Encombe valley is followed by cliffs rising to Hounstout, then descending to Chapman’s Pool, a rounded cove formed where the sea has scooped out soft Kimmeridge Clay between bouldery spurs at the mouth of a valley. The coast turns southward as cliffs run out to St. Alban’s (St Aldhelm’s) Head, (108 m), then eastward as a coastal slope descends to cliffs up to 60 m high in hard stratified Portland Stone towards Durlston Head. The more exposed southern shore of Durlston Head is boldly cliffed, but to

the north, in Durlston Bay, the east-facing coast consists of steep, crumbling slopes exposing the sequence of Purbeck Beds rising northward. Swanage Bay has been cut into the Wealden Beds at the eastern end of the lowland that extends through the Isle of Purbeck. At the northern end of Swanage Bay in Punfield Cove the steeply dipping Wealden Beds are overlain by Lower Greensand, Gault and Upper Greensand beneath the Ballard Down Chalk, which forms an escarpment cliff out to Ballard Point (>Fig. 7.12.11).

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⊡⊡ Fig. 7.12.11 Escarpment cliff in Chalk at Ballard Down. (Courtesy Geostudies.)

⊡⊡ Fig. 7.12.12 Chalk stacks at Handfast Point, on the south coast of Studland Bay. (Courtesy Geostudies.)

On the east-facing coast of Ballard Down the vertical cliffs decline to Handfast Point. In contrast to the similar coast at the eastern end of the Isle of Wight there is no shore platform. Towards Handfast Point the Chalk cliff becomes strongly crenulate, with bays and promontories cut along joint planes, erosion having separated an elongated peninsula into the stacks known as The Pinnacles and Old Harry Rocks (>Fig. 7.12.12). At the southern end of Studland Bay an eroded surface in Upper Chalk is overlain by the Reading Beds. Along the

shore are some large blocks of ferruginous sandstone and a beach of smooth, rounded chalk cobbles and pebbles and mainly unworn or broken flint nodules. The chalk cobbles and pebbles diminish rapidly northward and disappear as the result of abrasion (May and Heeps 1985). North of Redend Point the beach becomes sandy, backed by parallel dune ridges on the South Haven Pen­ insula. Beyond is Shell Bay, a curving NW to South Haven Point at the entrance to Poole Harbour, which has been constricted by formation of bordering spits, South Haven

Hampshire

Peninsula to the south and the Sandbanks peninsula extending SW from Bournemouth beach. Poole Harbour is a tidal estuary fed by several rivers, the largest of which are the Frome and the Piddle, entering at the western end. It shows the branched configuration typical of a valley system partly drowned by the Late Quaternary marine transgression, interfluvial ridges persisting as promontories and high islands. A century ago Poole Harbour had relatively narrow fringing salt marshes and wide soft mudflats exposed at low tide, but the ­salt-tolerant rice (cord) grass, Spartina anglica, became established in 1899 and quickly spread across intertidal areas to form wide swards (Bird and Ranwell 1964). The Spartina marshland is generally wider on shores sheltered from strong wave action, as on the eastern sides of islands and peninsulas, but thins out where tidal channels run close to the shore (May 2006). Sandbanks spit, formerly dune-capped, is now built over. East from Sandbanks former cliffs have been almost entirely stabilised as artificial vegetated slopes extending past Bournemouth to Branscombe. Cliffs then rise to Hengistbury Head. Offshore mining of ironstone nodules in the nineteenth century deepened the water and allowed stronger waves to attack the cliffs, which then began to retreat rapidly. In 1938 a groyne was built at the  end of the promontory, and this trapped eastward-­ drifting sand and gravel to form a widening beach, backed by low dunes in front of a cliff that became weathered back to a gentler bluff, with basal fans of sand and gravel no longer reached by the sea. The eastern cliff of Hengistbury Head then lost its beach, and was retreating with basal heaps of talus and ironstone, until recently, when it was stabilised behind an artificial beach held by several boulder groynes. A spit runs out northward to the entrance to Christchurch Harbour. Fed by the Rivers Stour and Avon,

7.13

Christchurch Harbour is a smaller version of Poole Harbour, with salt marshes, including some Spartina, bordering tidal mudflats and creeks. From Mudeford the coast is protected by sea walls, and the former cliffs in soft clay and sand have been largely stabilised as sloping vegetated bluffs. A series of massive groynes between sand and shingle beaches continues eastward in front of the stabilised coastal slopes, and this largely man-made coastal scenery runs on to the steep-sided chine at Chewton Bunny where the Dorset coast comes to an end at the Hampshire border.

References Arber MA (1973) Landslips near Lyme Regis. Proc Geol Assoc 84:121–133 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Bird ECF (1995) Geology and scenery of Dorset. Ex-Libris Press, Bradford on Avon, Wiltshire, pp 207 Bird ECF, Ranwell DS (1964) The physiography of Poole Harbour, Dorset. J Ecol 52:355–366 Brunsden D, Jones DKC (1972) The morphology of degraded landslide slopes, south-west Dorset. Quart J Engrg Geol 5:205–222 Brunsden D, Jones DKC (1976) The evolution of landslide slopes in Dorset. Philos Trans R Soc Lond A 283:605–631 Brunsden D, Jones DKC (1980) Relative time scales and formative events in coastal landslide systems. Zeitschrift für Geomorphologie, Supple­ ment band 34:1–19 Bury H (1920) The chines and cliffs of Bournemouth. Geol Mag 57:71–76 Carr AP, Blackley MWL (1974) Ideas on the origin and development of Chesil Beach, Dorset. Proc Dors Arch Soc 95:1–9 Davies KH, Keen DH (1985) The age of Pleistocene marine deposits at Portland, Dorset. Proc Geol Assoc 96:217–225 May VJ (2006) The ecology of Poole Harbour. Elsevier, Oxford May VJ, Heeps C (1985) The nature and rates of change on chalk coastlines. Zeitschrift für Geomorphologie, Supplement band 57:81–94 Wanklyn F (1922) History of Lyme Regis. Humpreys, London

7.13  Hampshire 1.  Introduction Hampshire occupies the northern part of a broad syncline that runs from west to east through The Solent and Spithead. The Chalk of the Dorset Downs and Salisbury Plain subsides beneath Tertiary formations comprising the Bagshot Beds (Poole Formation), overlain by the Lower Oligocene Barton and Bracklesham Beds and the Upper Oligocene Hamstead and Bembridge Beds: a succession of sands, clays, marls and some limestones. In

Late Tertiary times the so-called Solent River was an eastward extension of the Dorset Frome eastward through a valley which ran along the syncline, and the geological formations are sequences of marine, estuarine, deltaic and valley floor deposits that formed as sea level rose and fell (marine transgressions and regressions) along this valley (Everard 1954). The southern side of the valley was for long the Chalk ridge of the Purbeck Hills extending eastward through the centre of the Isle of Wight. This was breached by marine erosion to form the 24 km wide gap

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Peninsula to the south and the Sandbanks peninsula extending SW from Bournemouth beach. Poole Harbour is a tidal estuary fed by several rivers, the largest of which are the Frome and the Piddle, entering at the western end. It shows the branched configuration typical of a valley system partly drowned by the Late Quaternary marine transgression, interfluvial ridges persisting as promontories and high islands. A century ago Poole Harbour had relatively narrow fringing salt marshes and wide soft mudflats exposed at low tide, but the ­salt-tolerant rice (cord) grass, Spartina anglica, became established in 1899 and quickly spread across intertidal areas to form wide swards (Bird and Ranwell 1964). The Spartina marshland is generally wider on shores sheltered from strong wave action, as on the eastern sides of islands and peninsulas, but thins out where tidal channels run close to the shore (May 2006). Sandbanks spit, formerly dune-capped, is now built over. East from Sandbanks former cliffs have been almost entirely stabilised as artificial vegetated slopes extending past Bournemouth to Branscombe. Cliffs then rise to Hengistbury Head. Offshore mining of ironstone nodules in the nineteenth century deepened the water and allowed stronger waves to attack the cliffs, which then began to retreat rapidly. In 1938 a groyne was built at the  end of the promontory, and this trapped eastward-­ drifting sand and gravel to form a widening beach, backed by low dunes in front of a cliff that became weathered back to a gentler bluff, with basal fans of sand and gravel no longer reached by the sea. The eastern cliff of Hengistbury Head then lost its beach, and was retreating with basal heaps of talus and ironstone, until recently, when it was stabilised behind an artificial beach held by several boulder groynes. A spit runs out northward to the entrance to Christchurch Harbour. Fed by the Rivers Stour and Avon,

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Christchurch Harbour is a smaller version of Poole Harbour, with salt marshes, including some Spartina, bordering tidal mudflats and creeks. From Mudeford the coast is protected by sea walls, and the former cliffs in soft clay and sand have been largely stabilised as sloping vegetated bluffs. A series of massive groynes between sand and shingle beaches continues eastward in front of the stabilised coastal slopes, and this largely man-made coastal scenery runs on to the steep-sided chine at Chewton Bunny where the Dorset coast comes to an end at the Hampshire border.

References Arber MA (1973) Landslips near Lyme Regis. Proc Geol Assoc 84:121–133 Bird ECF (1989) The beaches of Lyme Bay. Proc Dors Arch Soc 111:91–97 Bird ECF (1995) Geology and scenery of Dorset. Ex-Libris Press, Bradford on Avon, Wiltshire, pp 207 Bird ECF, Ranwell DS (1964) The physiography of Poole Harbour, Dorset. J Ecol 52:355–366 Brunsden D, Jones DKC (1972) The morphology of degraded landslide slopes, south-west Dorset. Quart J Engrg Geol 5:205–222 Brunsden D, Jones DKC (1976) The evolution of landslide slopes in Dorset. Philos Trans R Soc Lond A 283:605–631 Brunsden D, Jones DKC (1980) Relative time scales and formative events in coastal landslide systems. Zeitschrift für Geomorphologie, Supple­ ment band 34:1–19 Bury H (1920) The chines and cliffs of Bournemouth. Geol Mag 57:71–76 Carr AP, Blackley MWL (1974) Ideas on the origin and development of Chesil Beach, Dorset. Proc Dors Arch Soc 95:1–9 Davies KH, Keen DH (1985) The age of Pleistocene marine deposits at Portland, Dorset. Proc Geol Assoc 96:217–225 May VJ (2006) The ecology of Poole Harbour. Elsevier, Oxford May VJ, Heeps C (1985) The nature and rates of change on chalk coastlines. Zeitschrift für Geomorphologie, Supplement band 57:81–94 Wanklyn F (1922) History of Lyme Regis. Humpreys, London

7.13  Hampshire 1.  Introduction Hampshire occupies the northern part of a broad syncline that runs from west to east through The Solent and Spithead. The Chalk of the Dorset Downs and Salisbury Plain subsides beneath Tertiary formations comprising the Bagshot Beds (Poole Formation), overlain by the Lower Oligocene Barton and Bracklesham Beds and the Upper Oligocene Hamstead and Bembridge Beds: a succession of sands, clays, marls and some limestones. In

Late Tertiary times the so-called Solent River was an eastward extension of the Dorset Frome eastward through a valley which ran along the syncline, and the geological formations are sequences of marine, estuarine, deltaic and valley floor deposits that formed as sea level rose and fell (marine transgressions and regressions) along this valley (Everard 1954). The southern side of the valley was for long the Chalk ridge of the Purbeck Hills extending eastward through the centre of the Isle of Wight. This was breached by marine erosion to form the 24 km wide gap

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Hampshire

now seen between Ballard Down in Dorset and The Needles at the western end of the Isle of Wight, and thus opens up Bournemouth Bay, Christchurch Bay and The Solent along the line of the submerged valley (Melville 1982). The coast of Hampshire is generally low-lying and fringed mainly by sand and shingle beaches and salt marshes. There are a number of estuarine inlets, as at Lymington and the Beaulieu River, and on a larger scale Southampton Water (the drowned mouths of the Test, Itchen and Hamble valleys), Portsmouth Harbour, Lang­stone Harbour and Chichester Harbour. Each of these is the outcome of marine submergence during the Late Quaternary (Flandrian) marine transgression, and they remain relatively open because the rivers draining into them have carried only small sediment loads. Mean spring tide ranges are generally small, with double high tides in the Solent, Southampton Water and Spithead. The range at Christchurch is 1.4 m and on Hurst Castle Spit 2.2 m. Lymington has 2.5 m, Calshot Castle 3.8 m and Southampton 4.0 m. At Lee on Solent the tides attain 3.7 m and at Portsmouth 4.1 m.

2.  The Hampshire Coast The coast of Christchurch Bay consists of steep bluffs and cliffs in the yellow and white Highcliffe Sands at the top of the Bracklesham Beds. At various times a spit grew

a­ longshore from Mudeford as far as Highcliffe Castle, deflecting the outlet from Christchurch Harbour eastward, but the harbour entrance has been stabilised by breakwaters. East of Highcliffe the Bracklesham Beds dip below sea level and an upper cliff of gravel-capped Barton Clay develops behind a slumped foreground, where mudflows and clay lobes have spread out across the shore (> Fig. 7.13.1). The cliffs rise to more than 35 m at Barton-on-Sea. There has been general recession of the cliffed coastline of Christchurch Bay during the past century. The crest of the upper cliff at Barton-on-Sea has retreated up to a metre per year, and several buildings have been lost. Sectors of the coastline have now been stabilised (> Fig. 7.13.2). The shingle beach continues along the coast, much of which is now modified by walls or wooden revetments. Near Milford on Sea headwater valleys of Dawes Stream have been beheaded by cliff recession to form cliff cols (similar to the valleuses of the Normandy coast). The cliffs decline east of Milford-on-Sea and the shingle beach continues as a high bank that ends in the recurved spit at Hurst Castle Point. In its lee are extensive salt marshes and mudflats exposed at low tide off Keyhaven, where the lowlying hinterland is protected by a high tide embankment at Sea Grass Lane. There is an intricate system of tidal creeks and some shelly cheniers on the salt marsh, which has an eroding microcliff (up to 1.5 m high) along its seaward margins (Pye and French 1993). Introduction of Spartina anglica late in the nineteenth century resulted in rapid vertical accretion and widening of the salt marsh terrace, but since 1930 there has been die-back of Spartina ⊡⊡ Fig. 7.13.1 Slumping clay cliffs at Barton, Christchurch Bay. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.13.2 Stabilised coast at Barton-on-Sea. (Courtesy Geostudies.)

(­ preceding erosion) along the seaward margins and beside tidal creeks, and in pans within the marshes (Bradbury 1996). Hurst Castle spit grew from the west, supplied with shingle drifting along the shore from Christchurch Bay (Lewis 1931; King and McCullach 1971). It has been shaped by SW waves, which have piled up Hurst Beach and driven it landward (over salt marsh) during storm surges (as in 1989–1990, when it moved up to 80 m), and easterly waves from the Solent, which have shaped the successive parallel recurves and built beach ridges at the eastern end (> Fig. 7.13.3). Narrow beaches back salt marshes along the coast east from the Lymington estuary, and there are successive spits at Warren Farm and Needs Oar (Ore) Point. At Inchmery Quay the salt marsh is being dissected by erosion (> Fig. 7.13.4). Brown flint shingle dominates Calshot spit, which curves out to Calshot Castle. It shelters salt marshes which extend NW along the shore of Southampton Water to Hythe. NW of Hythe former salt marshes have been ­extensively embanked and reclaimed for port and industrial development, but a cliffed salt marsh persists at Totton. There has been much land reclamation north of Southampton Water for quays and docks and for urban or industrial use, but south of the Itchen estuary, narrow sand and shingle beaches front low bluffs and cliffs in Bracklesham Beds at Netley and Warsash. Exposure to SW wave action increases towards Lee-on-the-Solent, where there has been cliff recession; but sea walls have been built to halt erosion here and in Stokes Bay. A broad

lobate shingle foreland has formed at Gilkicker Point, but the coast is now largely walled at Gosport and Portsmouth, with extensive docks. Within Portsmouth Harbour there are salt marshes and intertidal mudflats. Portsea Island is largely urbanised, with reclaimed areas along its coasts. The southern (Spithead) coast has Southsea promenade, and is fringed by a shingle beach (largely submerged at high tide) retained by a lattice pattern of groynes. The entrance to Langstone Harbour is bordered by inwashed paired shingle and sand spits. Glacial erratics have been found at the entrance to Langstone Harbour, possibly deposited from melting icebergs in Pleistocene times. Langstone Harbour is much like Portsmouth Harbour, with salt marshes, mudflats and tidal creeks converging southward, and much of the coastline walled at high tide level. The bordering beaches at Eastney and Hayling Island are mainly of brown and white subangular flint shingle, with white shelly berms. Hayling Island is low-lying and covered with brown Pleistocene brickearth, part of the wind-deposited loess that extends eastward along the Sussex coastal plain. There are minor bluffs, fringed by salt marshes and parallel shingle beach ridges on scrubby Sinah Common, in the SW of the island, which are surmounted by dunes. Salt marshes are extensive on the eastern shore. A shingle beach runs along the south coast of the island, with a sea wall and numerous groynes, to Eastoke Point. The entrance to Chichester Harbour is also bordered by paired spits, and the Sussex boundary runs down Emsworth Channel to the west of Thorney Island.

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⊡⊡ Fig. 7.13.3 Shingle beach ridges at the end of Hurst Castle spit. (Courtesy Geostudies.)

⊡⊡ Fig. 7.13.4 Eroding salt marsh at Inchmery Quay. (Courtesy Geostudies.)

References Bradbury A (1996) Western Solent salt marsh study. National Rivers Authority, Bristol, UK Everard CE (1954) The Solent river: a geomorphological study. Trans Inst Br Geogr 20:41–54 King CAM, McCullach MJ (1971) A simulation model of a complex recurved spit. J Geol 79:22–36

Lewis WV (1931) Effect of wave incidence on the configuration of a ­shingle beach. Geogr J 78:131–148 Melville RV (1982) The Hampshire Basin and adjoining areas. British Regional Geology, 4th edn. HMSO, London Pye K, French PW (1993) Erosion and accretion on British salt marshes. Cambridge Environmental Research Consultants, Cambridge, UK

Isle of Wight

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7.14  Isle of Wight 1.  Introduction The Isle of Wight is a diamond-shaped island measuring 37.8 km from east to west and 21.4 km from north to south, with a coastline about 90 km long. Its central feature is a ridge of steeply-dipping Chalk extending from The Needles in the west to Culver Cliff in the east. To the north are Tertiary formations, Eocene and Oligocene: the Eocene strata stand almost vertical in Alum Bay and Whitecliff Bay, immediately adjacent to the Chalk ridge, but the dip diminishes rapidly northward and the Upper Oligocene strata are almost horizontal. To the south the Chalk escarpment looks across the outcrops of Lower Cretaceous formations exposed in anticlinal zones to a southern plateau of Chalk dipping gently southward (Bird 1997). Much of the coast is cliffed across these various formations. The drainage divide is generally close to the southern coast, so that relatively long rivers drain to the north coast estuaries, as at Newtown Harbour, the Medina at Cowes, Wootton Creek and Bembridge Harbour. Shorter streams drain to the south coast, several of them having cut deep, narrow steep-sided valleys known as chines. Beaches are of sand and shingle, the proportion of shingle increasing in the vicinity of sandy cliffs, and salt marshes are found in the estuaries. Mean spring tide range is small around the Isle of Wight, where double high tides occur on the Solent and Spithead shores. The range at Cowes is 3.6 m, diminishing to 2.6 m at Yarmouth and 2.2 m in Totland Bay on the west coast. On the south coast the tide range at  Freshwater is 2.1 m, increasing to 3.1 m at Ventnor and 3.5 m at Sandown. On the east coast Brading Harbour has 3.1 m, and the largest tides are at Ryde, 3.8 m. The coast is described and illustrated in an anticlockwise sequence beginning and ending at Cowes in the north.

2.  The Isle of Wight Coastline The Medina estuary opens northward from a steep-sided, almost straight drowned valley north from Newtown to open to the sea beside Cowes on its west bank. Cowes stands on a hill where slopes in Bembridge Marls descend to Bembridge Limestone on the shore. The promenade, with a sea wall and a narrow beach of flint gravel, curves round to Egypt Point. To the west vegetated coastal bluffs rise behind Gurnard Bay. The coastal slope is a degraded

escarpment, with hummocky slopes on the Bembridge Marls and patchy outcrops of Bembridge Limestone along the shore. In Thorness Bay the wide muddy foreshore exposed at low tide has ledges of Bembridge Limestone, and is backed by sandy beach ridges on a spit that deflects the little stream eastward. Cliffs rise to the west and the beach becomes a shingle spit, capped by dunes, on the western side of Newtown Harbour. This is a ria formed by Late Quaternary marine submergence, which was followed by the growth of paired spits bordering the marine entrance. West of Newtown Harbour, cliffs and bluffs extend to Yarmouth, where the Western Yar has an estuary bordered by tidal mudflats and salt marshes with areas of Spartina grass, threaded by tidal creeks. At Sconce Point the coast turns southward, and there are cliffs cut in almost horizontal marls and limestones, past Totland Bay and beneath Headon Hill. In the northern part of Alum Bay the beach is sandy and the cliffs show chutes of slumping clay and lobes of truncated Headon Beds. The strata curve upward and are soon almost vertical as the pale yellow and grey Barton Sand emerges north of Alum Bay Chine. There follows a descending sequence of Eocene and Palaeocne strata standing almost vertical in the cliffs. The cliff face has been dissected into ravines on soft strata and ribs on harder rock in the almost vertical strata. At the southern end the basal Reading Beds dip northward at about 80°, immediately overlying eroded Upper Chalk, which emerges as steep cliffs along the southern coast of Alum Bay. The Upper Chalk cliffs run out westward, and have grassy slope segments and buttresses and a narrow basal ledge. A rock fall in 1993 produced a white scar (an exposed nearly vertical bedding plane), with a large talus apron, which was vegetated and undercut by a basal cliff in 2006: the scar was then grey (> Fig.  7.14.1). The grassy slope has been cut back from the plane of the sub-Palaeocene surface, and is not strictly a dip-slope. The Needles are stacks of very hard and splintery Upper Chalk which has become resistant largely as the result of compression and recrystallisation accompanying the intense folding. Three large stacks running out to the red and white lighthouse are shards of this chalk, about 30 m high (>Fig. 7.14.2). The Chalk cliffs run behind Scratchell’s Bay, where the northward-dipping Upper Chalk is crossed by numerous steeply inclined parallel flint horizons marking the ­bedding-planes. There is a beach of blue-black flint shingle  eroded from the cliffs, which undulate and decline ­westward to Watcombe Bay, below the steep convex grassy

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⊡⊡ Fig. 7.14.1 Slumped fan of Chalk undercut by marine erosion on the south coast of Alum Bay. (Courtesy Geostudies.)

escarpment of Tennyson Down. Watcombe Bay is a small cove cut into the head of a dry valley, and gravelly Coombe Rock in the cliff indicates that a deeper valley was excavated, then infilled with frost-shattered chalk and flint sludged down the bordering slopes under periglacial conditions in the Late Pleistocene. Cliffs cut in hard Upper Chalk descend into Freshwater Bay, a broad cove at the head of the Western Yar valley. It is similar to Lulworth Cove in Dorset, but with a wider entrance, and cut entirely in Chalk. The eastern half of the cove has a beach of flint shingle in front of cliffs cut in shattered Upper Chalk, the flint layers having been much disturbed by frost heaving and lateral sludging. This is overlain by Coombe Rock and a reddish brown Brickearth with some interbedded gravel. There are residual stacks, Stag Rock, Mermaid Rock and the foundations of Arch Rock, on the shore platform below the cliff to the east. The Chalk cliffs continue into Compton Bay, passing into the red ferruginous sands of the Upper Greensand (> Fig. 7.14.3). The shore is littered with chalk and greensand boulders. The Upper Greensand recedes behind a landslide in Atherfield Clay as the coast curves out to Hanover Point. The SW coast of the Isle of Wight has cliffs cut in Lower Cretaceous (Wealden) marls and clays with beds of sandstone forming shore ledges. They are incised by several small V-shaped chines, notably Alum Chine to the SE. An undercliff terrace develops near the cliff crest on the eastern side of Walpen Chine and widens to 130 m as it

descends eastward. Springs and seepages of groundwater from thick overlying sandstones issue from the cliff base, especially after wet weather, and this outflow has been res­ ponsible for the undermining, collapse and recession of the sandstone cliffs, as well as the consequent disappearance of Blackgang Chine. Cliffs in sandrock and sandstone stand behind a subsided undercliff (> Fig.  7.14.4) and recurrent landslides have damaged and destroyed cottages and chalets at Blackgang (Hutchinson et al. 1981). To the east the high vertical upper cliff swings seaward and the undercliff terrace comes to an end. Almost verti­­cal  cliffs in Upper Greensand extend past Rocken End. Above Rocken End the irregular topography of the Niton Landslide (> Fig. 7.14.5) ascends to the base of Gore Cliff, a vertical upper cliff in evenly-bedded horizontally stratified Upper Greensand, with protruding ledges of hard nodular cornstones and etched out shelves in softer sandstone. It is capped by the Lower Chalk, which forms a receded brow behind a wind-scoured bench. Recession of the upper cliff has interrupted the old undercliff road at Windy Corner. The steep coast east towards Ventnor consists of wooded slopes on talus and slumped rock, with basal cliffs cut into slipped masses of Chalk and Upper Greensand, often with a landward dip, indicating rotational slumping, extending past St. Catherine’s Point (Hutchinson et al. 1991). Ventnor originated as a seaside resort and spa because of a spring near the base of the slumped cliffs. Sea walls of varying design have been built eastward to Horseshoe Bay,

Isle of Wight

⊡⊡ Fig. 7.14.2 Scratchells Bay, looking towards The Needles. The sea is milky with chalk powder in suspension. (Courtesy Geostudies.)

⊡⊡ Fig. 7.14.3 Steeply dipping Chalk underlain by Upper Greensand, Gault and Lower Greensand (Ferruginous Sands) in Compton Bay. (Courtesy Geostudies.)

7.14

fronting cliffs in slumped material, some of which have been graded to grassy slopes. The Undercliff to the east is a broad strip of descending terraces of Chalk and Upper Greensand that have subsided over the Gault, overlooked by an upper cliff in which layers of Upper Greensand outcrop, were overlain by the Lower Chalk. There is a wide perched basin behind Luccombe Bay floored with Gault, and overlooked by an escarpment of Upper Greensand and Lower Chalk. Luccombe Chine is a deeply incised wooded ravine cut through the lower cliff of Sandrock Beds, and above Shanklin Point is bold Knock Cliff - the lower part showing dark Ferruginous Sands, the upper part massive grey Sandrock Beds capped by yellow bedded sandstones (> Fig. 7.14.6). From the top of Knock Cliff the ground drops across to hummocky, slumped topography on the Gault, and the cliffs decline to Appley Beach at the western end of Shanklin. The cliff, draped with vegetation, runs behind the Chair Lift at Shanklin, where the Ferruginous Sands show a gentle northward dip on the southern side of the gentle Shanklin Anticline. The coast between Shanklin and Sandown consists of almost vertical cliffs 20–30 m high in Ferruginous Sands, with basal talus covered by grasses and scrub, fronted by a sea wall and undercliff walk. Sandown is built on a plateau of Ferruginous Sands, and the coast curves gently round to the east. Sandown Bay has been excavated in the Wealden Beds which occupy the core of the unroofed Sandown Anticline. Beyond the last of the many wooden groynes at San­ down, built to prevent the sand drifting away eastward,

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⊡⊡ Fig. 7.14.4 Blackgang Undercliff. (Courtesy Geostudies.)

⊡⊡ Fig. 7.14.5 Receding cliffs in Upper Greensand above the landslide at Rocken End. (Courtesy Geostudies.)

slumping cliffs in Wealden Beds and Atherfield Clay pass into vertical cliffs in the more resistant Ferruginous Sands and the Sandrock Beds. The cliff crest declines through a hollow across the head of a dry valley that descends to Yaverland. There is a hollow on a narrow outcrop of Gault clay, followed by the Upper Greensand, somewhat ob­­ scured by slumping, and then the Lower Chalk descending from the cliff crest. Culver Cliff has been cut into the south-facing Chalk escarpment by marine erosion. At the eastern end, Whi­ tecliff Point is a vertical cliff in hard, massive Upper Chalk,

with irregular flint layers including some large tabular nodules along the bedding planes, and seams of marl. At the southern end is the jutting spur called the Anvil, formed of massive white chalk with few marl seams, its grassy northern slope following a bedding-plane which has been exposed by erosion along a weaker overlying horizon. The shore platform is strewn with chalk bould­­ ers, and shows steeply-inclined ribs that extend eastward between wave-scoured furrows, some with elongated pools at low tide. There are protruding layers and nodules of flint.

Isle of Wight

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⊡⊡ Fig. 7.14.6 Knock Cliff. (Courtesy Geostudies.)

On the southern side of Whitecliff Bay the cliff is cut into Upper Chalk, high and partly vegetated with talus slopes. It is very steep, with a large notch at the base, and a shore platform, cut across steeply dipping chalk strata and submerged at high tide. This part of the coast is much less exposed to storm waves than Alum Bay and the top of the Chalk is obscured by grassy slumped aprons. The Eocene formations in Whitecliff Bay are similar to those in Alum Bay at the western end of the Isle of Wight, showing the same sequence of vertical strata, the dip declining to the north. At the northern end of Whitecliff Bay the Bembridge Limestone appears in the top of the cliff and descends northward. Low cliffs and ledges of brown Bembridge Limestone continue along the shore, the inner and outer reefs separated by a corridor of gravelly sand which declines into a low tide channel known as The Run. Off The Foreland, the Bembridge Ledges are parallel outcrops of Bembridge Limestone. Low slumping cliffs cut in dark Bembridge Marls, capped by gravels in a clay matrix, extend NW past Bembridge. Bembridge Point is a broad, blunt dune-capped spit with a shingle beach fronted by sand on its seaward side. The Yar estuary which has been formed by submer­­gence  of a river valley cut out along the eastern part of the Bouldnor syncline widens into a bay - Bembridge Harbour. To the north the Bembridge Limestone reappears in a series of rocky spurs, between which are little beaches of sand and flint pebbles. An outcrop of hard calcareous sandstone is

responsible for the Horestone Point promontory, which declines to a shore where the Nettlestone Grit has disintegrated into large reddish-brown and ­ferruginous blocks. At Seaview shore ledges are on the outcrop of this grit, which at Nettlestone Point can be traced laterally from a shelly limestone into a hard marl, and then calcareous sandstones with flint pebbles. The coast curves out to Puckpool Point, and there are walled sectors westward to Ryde. An extensive sandy area is exposed at low tide in front of the Ryde promenade. Much of the coast between Ryde and Cowes consist of low wooded bluffs with minor basal cliffing and narrow gravel beaches, interrupted by the estuary of Wootton Creek, another drowned valley, with paired spits bordering the entrance. At Old Castle Point the sea wall and coast road begin, curving round to the mouth of the Medina estuary at Cowes.

References Bird ECF (1997) The Shaping of the Isle of Wight. Ex-Libris Press, Bradford on Avon, Wiltshire Hutchinson JN, Chandler MP, Bromhead EN (1981) Cliff recession on the Isle of Wight south-west coast. In: Proceedings of the 10th internatio­ ­nal  conference on soil mechanics and foundation engineering, Stockholm, Sweden, pp 429–434 Hutchinson JN, Brunsden D, Lee EM (1991) The geomorphology of the landslide complex at Ventnor, Isle of Wight. In: Chandler RJ (ed) Slope stability engineering. Thomas Telford, London, pp 213–218

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7.15  Sussex 1.  Introduction The western part of the Sussex coast is low-lying, bordering a coastal plain from the Hampshire border to Brighton, where it is intersected by the South Downs. Chalk cliffs extend from Brighton to Eastbourne interrupted by gaps at the mouths of the Ouse and Cuckmere valleys, and beyond Eastbourne the coast cuts across the southern part of the Wealden dome. Pevensey Levels are alluvial lowlands at the eastern end of the Sussex Low Weald (the Weald Clay vale), and were swamps and fens until they were embanked and reclaimed. At Bexhill, the Tunbridge Wells Sandstone reaches the coast in low cliffs and shore reefs, and beyond Hastings the high sandstone cliffs are cut in Ashdown Sand and Fairlight Clay in the central anticline of the High Weald. Another alluvial lowland at Winchelsea and Rye passes behind the large shingle complex at Dungeness Foreland. Mean spring tide range on the Sussex coast increases gradually eastward from 4.2 m at the entrance to Chichester Harbour to 4.7 m at Selsey Bill, 4.9 m at Pagham, 5.0 m at Bognor Regis, 5.2 m at Littlehampton, 5.4 m at Worthing, 5.5 m at Shoreham, 5.9 m at Brighton, 6.1 m at Newhaven, 6.5 m at Eastbourne and 6.8 m at Hastings.

2.  The Sussex Coast The Hampshire-Sussex boundary runs down Emsworth Channel in Chichester Harbour, to the west of Thorney

Island, which has been attached to the mainland by land reclamation. It is bordered by sea walls overlooking salt marshes and mudflats exposed at low tide (> Fig.  7.15.1), the salt marshes including much Spartina. On the southern shore the salt marsh is much dissected and eroding, and there are some large boulders of distant origin, thought to have been deposited from melting icebergs Late Pleistocene times. The streams draining into Chichester Harbour carry very little sediment, and the mud that has accreted has been derived largely from the brickearth deposits on the coastal plain. Paired spits border the entrance to Chichester Harbour, the one on the eastern (Sussex) side being East Head, a recurved sand and shingle spit surmounted by dunes up to 4 m high and backed by salt marshes and a shallow bay at high tide. The Chidham peninsula is similar to Thorney Island, and on its eastern side Bosham Channel is a tributary of the large Chichester Channel. At Bosham high tide flooding is frequent, especially at the equinoxes (> Fig. 7.15.2). East of the entrance to Chichester Harbour the coast of Bracklesham Bay has a shingle beach and numerous groynes. Wave action is limited by the short fetch across Spithead, but there are some low vertical cliffs in Brickearth at West Wittering. At Selsey Bill erosion has been rapid, but sectors are now protected by boulder ramparts. The shingle beach continues northward to the paired spits that border the entrance to Pagham Harbour. This bay has extensive salt marshes and mudflats at low tide, and receives only minor freshwater inflow from streams and drainage ditches. ⊡⊡ Fig. 7.15.1 Dissected salt marsh, Thorney Island. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.15.2 High spring tide flooding at Bosham. (Courtesy Geostudies.)

Gravelly shoals are extensive off Pagham Harbour, and to the east sandstone reefs outcrop offshore at Barn Rocks and Bognor Rocks. The seafront at Bognor Regis has been walled and armoured with dumped boulders, and numerous groynes have been built to prevent beach shingle drifting away to the east. There has been artificial replenishment of the shingle beaches at Felpham and Middleton- on- Sea, where nearshore breakwaters have been constructed. At Atherington (Climping Beach) dunes back the shingle beach and a sandy foreshore is exposed at low tide, with green weed on gravel patches and in gravelly pools. The dunes continue east to the Arun mouth, where the beach widens because of accretion alongside the harbour breakwaters. East from Littlehampton the shingle beaches continue, with many groynes, along the edge of the urbanised Sussex coastal plain to Worthing. At Shoreham-by-Sea a shingle bank, built by eastward drifting, has forced the Adur outlet to the east. The beach at Hove has been depleted as the result of interception of drifting shingle by the breakwaters at the Adur mouth. The seafront has been walled from Hove past Brighton to Black Rock, and the shingle beach retained by numerous groynes. At Black Rock the Chalk cliffs begin. There is a section showing an emerged shingle beach overlain by frost-­ shattered chalk and flint gravel (Coombe Rock) behind the Black Rock marina, built in 1974. Recession of the Chalk cliffs to the east was rapid until an undercliff walk was built in the 1930s between Black Rock and Saltdean, serving as a high tide wall. The undercliff walk was subsequently extended eastward from Saltdean, but the Chalk

cliff is still receding at Peacehaven. It is fronted by a slightly concave shore platform up to 150 m wide, with transverse gradient declining seaward below low spring tide level. This shore platform has been cut across Chalk strata that dip gently seaward, forming minor scarps and mesas, particularly on flint layers and dissected into chalk boulders. It has been cut partly by abrasion, as waves move flint fragments to and fro and partly by solution of the Chalk in rainwater and seaspray: there are often clouds of fine calcium carbonate in suspension offshore, possibly formed as dissolved chalk is precipitated. Frost shattering during cold winters has contributed to weathering on the shore platform (Robinson and Jerwood 1987) (> Fig. 7.15.3). Shingle drifting eastward has accumulated on the beach that widens towards the Newhaven breakwater (> Fig. 7.15.4) built in 1731 and later extended. Behind the harbour, the Chalk is overlain by the Reading Beds at Castle Hill, and the former cliff, protected by the breakwater and reclaimed land, has developed a distinctive profile: an upper, convex slope on soft slumping Eocene sediment, a cliff face of vertical Chalk and a lower concave slope of downwashed chalk, sand and clay. The shingle beach continues to Seaford, where it was much depleted as a consequence of the interception of eastward drifting by Newhaven breakwater. Groynes failed to retain it, and a sea wall along the promenade had to be repaired after storms. In 1987 a shingle beach was renourished here. At the eastern end of the promenade frost-shattered Chalk rises in cliffs to Seaford Head (85 m). These diminish on the eastern side towards Cuckmere Haven, at the

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⊡⊡ Fig. 7.15.3 Shingle accumulation west of the Newhaven breakwater. The white strip is scoured Chalk exposed after waves pushed back the base of the beach. (Courtesy Geostudies.)

⊡⊡ Fig. 7.15.4 Chalk cliffs and intertidal shore platform, looking east towards Beachy Head. (Courtesy Geostudies.)

mouth of a wide valley cut through the South Downs. Beyond Cuckmere Haven the Seven Sisters are Chalk ridges and dry valleys truncated by vertical cliffs up to 70 m high which have been dissected along joints and flint-lined bedding planes. The Chalk has a very gentle southward dip, truncated in broad fronting intertidal shore platforms (> Fig. 7.15.4). The cliff face has a series of prominent buttresses (> Fig. 7.15.5). At Birling Gap the 12 m high cliffs show solid Chalk on either side of up to 5 m of Coombe Rock over

f­ rost-shattered Chalk beneath a broad dry valley. Cliff recession has truncated the row of nineteenth century fisherman’s cottages, and the steep stairway down to the beach has been repeatedly replaced. To the east the cliffs rise to Beachy Head (160 m), where the escarpment of the South Downs is truncated by the coastline. Cliff recession here is by way of successive rock falls, producing aprons of Chalk rubble that are gradually consumed by abrasion and solution. A major rock fall occurred in January 1999, producing a talus fan that extended out to

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⊡⊡ Fig. 7.15.5 The Chalk cliffs of the Seven Sisters, showing prominent vertical buttresses behind an intertidal shore platform. (Courtesy Geostudies.)

the lighthouse, and has since been trimmed back several metres (Williams et al. 2004). From Beachy Head the Chalk escarpment swings inland to face eastward, and the cliffs decline from the Upper Chalk of the escarpment crest through Middle and Lower Chalk to Upper Greensand and Gault, fronted by shingle and boulders at Holywell. Sea walls line the Eastbourne waterfront, with shingle beaches contained between groynes. To the east the shingle widens to ­multiple beach ridges on a blunt lobate foreland - The Crumbles out to Langney Point. Between Eastbourne and Hastings numerous groynes retain shingle within beach compartments, often angled in response to longshore drifting and the dominant SW waves. The shingle bank is backed by the wide lowlands of Pevensey Levels, former marshes reclaimed as pastureland. At Bexhill there are low cliffs and shore reefs cut into the Tunbridge Wells Sand Formation, and the shingle beach continues along the promenades of St Leonard’s and Hastings. From Hastings past Fairlight to Cliff End are high cliffs in the Hastings Beds of the Wealden Series, exposed in the core of the Wealden dome, where the High Weald comes to the coast. Cliffs cut in the pale brown, stratified, locally massive (sandrock) Ashdown Sand Formation pass into landslides as the underlying Fairlight Clay rises in the anticlinal ridge, which has been transected by marine erosion. The sandstone cliffs show sloping vegetated segments between vertical walls of more massive sandrock. Blocks of sandstone disintegrate along joints and bedding planes

and fall to the shore, where they form talus aprons and mingle with the eastward drifting- mainly flint shingle. Ecclesbourne Glen (> Fig. 7.15.6) and Fairlight Glen are deep valleys (known as ghylls in the High Weald) that have been truncated by recession of cliffs cut into the Ashdown Sand Formation, so that the streams end in waterfalls. East of Fairlight Glen the cliffs have collapsed into landslide lobes where the Fairlight Clay underlies massive Ashdown Sandrock. These lobes have segregated beach compartments and interrupted eastward drifting of beach shingle, so that the beach widens on their western flanks. Cliff recession has undermined houses at Fairlight Cove. At Cliff End (> Fig. 7.15.7) the cliffs of Ashdown Sand terminate and pass inland as bluffs that run behind Pett Level to Winchelsea. These were formerly coastal cliffs, but the longshore growth of a shingle spit resulted in the ­formation of marshes that have been drained and reclaimed as pastureland. The shingle spit has ridges diverging towards the River Rother. The Rother estuary has sloping mudbanks on either side of the low tide channel and breakwaters have been built to define its mouth. These have intercepted shingle drifting alongshore to form a wide shingle beach while to the east is a sandy foreshore (Camber Sands), backed by dunes. Camber Sands grade laterally into Broomhill Sands as the coast swings SE; the dunes come to an end and a shingle beach borders the sea wall that protects the highway as far as the mouth of a drain, Jury’s Gut. The shingle beach continues along the southern shore of Dungeness, across the Kent boundary.

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⊡⊡ Fig. 7.15.6 Cliffs east of Ecclesbourne Glen, showing capping sandrock above stratified Ashdown Sand. (Courtesy Geostudies.)

⊡⊡ Fig. 7.15.7 Cliffs in the Hastings Beds sandstones terminate at Cliff End, at the start of a shingle beach. (Courtesy Geostudies.)

References Robinson DA, Jerwood LC (1987) Frost and salt weathering of chalk shore platforms near Brighton, Sussex. Trans Inst Br Geogr 12:217–226

Williams RBG, Robinson DA, Dornbusch U, Foote YLM, Moses CA, Saddleton PR (2004) A sturzstrom-like cliff fall on the chalk coast of Sussex. In: Mortimore RN, Duperret A (eds) Coastal chalk cliff instability. Geological Society Engineering Geology Special Publication 20, pp 89–97

Kent

7.16

7.16  Kent The Kent coast begins on the south of the major cuspate foreland at Dungeness and extends round that foreland to Hythe, where it intersects geological formations of the northern part of the Wealden dome. The Lower Greensand escarpment runs WNW–ESE and comes close to the coast in a section that provided the type areas for the Hythe Beds, Sandgate Beds and Folkestone Sands. The parallel vale of the Gault then passes through Folkestone, and the Upper Greensand and Chalk escarpment also runs WNW–ESE above the major landslide at Folkestone Warren, reaching the coast in cliffs west of Dover. The Chalk cliffs extend past Dover and round to Walmer, where they decline to a synclinal lowland on Palaeocene rocks (Thanet Beds), the coast being fringed by dunes and Pleistocene Brickearth, seen in cliffs in Pegwell Bay. The Chalk rises again to the Isle of Thanet, forming cliffs from the northern side of Pegwell Bay round past Ramsgate and North Foreland to Margate and Westgate-on-Sea. The north-­facing coast then passes on to Lower Tertiary formations including Thanet Beds and London Clay. These extend past Herne Bay and Whitstable, dominate the northern part of the Isle of Sheppey and extend through the Hoo Peninsula north of the Medway estuary. Much of the coast is lowlying land that has been embanked, drained and reclaimed, as in Graveney Marshes and Luddenham Marshes south of The Swale, the channel that isolates the Isle of Sheppey, with Eastchurch, Elmley and Minster ⊡⊡ Fig. 7.16.1 Shingle beach ridges on Dungeness. (Courtesy Geostudies.)

Marshes on the southern part of that island, and the Isle of Grain. As the Thames estuary narrows upstream its shores are largely artificial banks and walls, although there is Chalk beneath the estuary and alongside it at Erith. Mean spring tide ranges diminish along the Kent coast from 7.0 m at Dungeness to 6.4 m at Folkestone, 5.9 m at Dover, 5.3 m at Deal, 4.5 m at Ramsgate, 4.2 m at Broad­ stairs and 4.3 m at Margate. There is then an increase ­westward to 4.4 m at Herne Bay, 4.9 m at Whitstable, 5.1 m at Sheerness, 5.2 m at Bee Ness, 5.4 m at Upnor and 5.6 m at Chatham. In the Thames estuary the tide range in­creases to 6.6 m at London Bridge.

The Kentish Coast Dungeness is a large cuspate foreland which in recent centuries has been shaped by the erosion of beach material from its southern shore, where earlier beach ridges have been truncated, longshore drifting past the rounded point and deposition on the eastern shore (Lewis 1932). Pro­gradation has continued on the eastern shore, with the addition of parallel shingle beach ridges and swales (>Fig.  7.16.1) up to Littlestone-onSea (New Romney). At Dymchurch a high sea wall (>Fig. 7.16.2), repaired and rebuilt after many storms, protects the lowland, and is

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Kent

⊡⊡ Fig. 7.16.2 The sea wall at Dymchurch. (Courtesy Geostudies.)

f­ ron­ted by a shingle beach descending to a wide sandy shore exposed at low tide. Much of the coastline east of Lympne is also fringed by sea walls, and groynes have been inserted in an attempt to stop shingle drifting away to the east. At Folkestone the beach widens towards the breakwater on the western side of the harbour, and beyond this the coast curves past East Cliff, a headland of Lower Greensand (Folkestone Sands), to Copt Point, bordering East Wear Bay. Folkestone Warren (>Fig. 7.16.3) is an area of recurrent landslides of tumbled and tilted rock where the Chalk and  Upper Greensand have subsided over the Gault Clay. The London to Dover railway was built through this area in 1844, and was frequently disrupted by ­landslides until massive concrete aprons were built (Hutchinson 1968). The landslide is backed by an upper cliff of Chalk that descends eastward to become a coastal cliff, part of which is protected by Samphire Hoe (>Fig. 7.16.4), an artificial terrace formed by land reclamation in front of the coast between Abbot’s Cliff and Shakespeare Cliff. Between 1990 and 1994 chalky waste excavated from the Channel Tunnel was dumped here. Shakespeare Cliff is an escarpment cliff where the Chalk dips northward. Dover Harbour began as an inlet at the mouth of a shingle-encumbered river, but the harbour was eventually enclosed by three large breakwaters. East of Dover the cliffs resume beneath Dover Castle and extend past South Foreland to St. Margaret’s Bay (>Fig.  7.16.5) and then north to Kingsdown. The Chalk is stratified and jointed, and forms vertical cliffs that recede intermittently

as the result of rock falls, especially during frosts when the rock is disrupted by expansion and contraction accompanying the freezing and thawing of interstitial water. The rock falls produce talus fans that are gradually consumed by the sea. The pattern of cliff recession is related to jointing (Middlemiss 1983). The Chalk cliffs drop to a shore platform up to 150 m wide when it is exposed at low spring tide, and traversed by grooves that run out at right angles to the coastline. The shingle beach is sparse, longshore drifting of shingle having been impeded by the breakwaters at Dover Harbour. At Kingsdown the shore platform gives place to a shingle beach which extends northward past Deal. Sandwich stands beside the estuary of the River Stour, which flows southward to round an inner spit at Richborough, then northward behind a broad outer spit which formed subsequently, before opening into Pegwell Bay. To the north low cliffs are cut in Pleistocene brickearth, a form of loess (aeolian silt) deposited under cold climatic conditions which is soft but coherent, forming vertical outcrops. Pegwell Bay has wide intertidal sand and mudflats, with a dissected Spartina marsh, backed by a sand and shingle beach. In the cliffs the brickearth is underlain by Palaeocene Thanet Sand over frost-shattered Chalk rubble. To the east along the south coast of the Isle of Thanet the base of the cliff is in solid Chalk, and abrasion has formed a slight notch fronted by an abrasion ramp that widens eastward into a shore platform. Caves have been cut out along fault lines (>Fig. 7.16.6). At Ramsgate a sandy beach develops over the back of the shore platform, but fades out northward. Sea walls

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⊡⊡ Fig. 7.16.3 Folkestone Warren. (Courtesy Geostudies.)

⊡⊡ Fig. 7.16.4 Samphire Hoe, an artificial terrace fronting Chalk cliffs west of Dover. (Courtesy Geostudies.)

begin at West Cliff and Ramsgate Harbour is enclosed by large curving breakwaters. The shore platform shows much karstic dissection, with Chalk pitted by rain and spray. Chalk cliffs continue northward past Dumpton Gap, where a sanded area interrupts the Chalk shore platform opposite a dry valley underlain by chalky rubble (Coombe Rock) to Broadstairs. High cliffs in solid Chalk then continue past North Foreland to Foreness Point, where the coast turns westward. An undercliff wall begins on Long Nose Point and runs to Cliftonville. The cliffs then decline

towards Margate, where the sandy beach is wide at low tide and the Chalk disappears near the ruins of Margate Pier. West of Margate the Chalk cliffs are low, and several sectors are walled. At Birchington there is a wide concrete undercliff walk in front of the former crenulate cliffs, which are vertical to overhanging and 5–6 m high. There are numerous small coves and headlands, the exposed Chalk having many walled-up caves and clefts. The Chalk shore platform is very wide at low tide.

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⊡⊡ Fig. 7.16.5 Chalk cliffs at St. Margaret’s Bay, east of Dover. (Courtesy Geostudies.)

⊡⊡ Fig. 7.16.6 Chalk cliffs and caves on the north coast of Pegwell Bay. (Courtesy Geostudies.)

At Minnis Bay, west of Birchington, the Chalk of the Isle of Thanet comes to an end, and the shore platform gives place to sandy mud on a wide intertidal shore. A sea wall protects reclaimed marshland in the former Wantsum Channel, which in Roman times was a strait 2 miles wide and up to 40 feet deep through to the Stour at Richborough on the east coast. A shingle beach runs in front of the sea wall to Reculver, and to the west are cliffs in Thanet Sand, then Woolwich and Reading Beds. Landslides and mudflows have occurred on coastal slopes where there has been

long-term cliff recession. At Beltinge former cliffs cut in London Clay have been landscaped as grassy bluffs with gravel drains, a basal sea wall, and groynes between sand and shingle beaches. The sea wall becomes the Herne Bay  promenade, and artificial grassy slopes continue to Hampton. Whitstable has a sea wall fronted by a shingle beach within compartments between wooden groynes. Bluffs that were formerly cliffs cut in London Clay extend eastward as the coast curves round to Seasalter. The southern coast of the Isle of Sheppey is artificial, with a sea wall protecting the reclaimed Minster, Elmley

Essex

and Eastchurch Marshes. At Shell Ness, the easternmost point, shelly beach sand drifting south–east has formed a small spit. The NE coast of the Isle of Sheppey rises to bluffs in London Clay at Leysdown on Sea and a prominent cliff, also in London Clay, on Warden Point. The coast to the west consists of eroding cliffs subject to recurrent landslides and mudflows on London Clay. Sheerness, on the NW coast, bordering the mouth of the Medway estuary, has a walled waterfront. Salt marshes in the Medway estuary are undergoing dissection and shrinkage, with marginal cliffing, the tidal creeks widening and intervening areas breaking up into grassy islets. At Allhallows, a ridge of London Clay is intersected by the southern shore of the Thames estuary, where the mudflats, over a mile wide at low tide, are backed by a thin shelly sand and pebble beach and a cliff cut in ­alluvium. Upstream the Thames swings southward to Shornmead Fort, then westward to Gravesend, and the estuary shore is bordered by artificial banks of earth or

7.17

concrete protecting low-lying reclaimed land that was formerly salt marsh or reed fen. Upstream are port and industrial installations, and the  Thames Barrier was built to exclude storm-driven high  tides from the Port of London. This Barrier has increased the risk of marine flooding downstream, leading  to the building of such structures as the Dartford Gate  at the mouth of the tributary Darent River, and surge gates on the promenades of bordering towns, as at Queenborough.

References Hutchinson JN (1968) The coastal landslides of Kent. Proc Geol Assoc 79:227–237 Lewis WV (1932) The formation of Dungeness foreland. Geogr J 80:309–324 Middlemiss FA (1983) Instability of chalk cliffs between the South Foreland and Kingsdown, Kent. Proc Geol Assoc 94:115–122

7.17  Essex 1.  Introduction Much of the Essex coast is low-lying and estuarine, and there has been extensive embanking and reclamation of former marshland. Salt marshes persist on the shores of estuaries, and on parts of the outer coast where there is a wide intertidal zone, but on the more exposed sectors there are beaches of sand, sometimes with shingle or shells. There are only a few cliffs, mainly cut in glacial drift, and some of these have been stabilised as bluffs behind sea walls, particularly at seaside resorts such as Clacton. North of the Thames estuary there are several long estuarine inlets, some of which may be inherited from former mouths of the Thames. Mean spring tide range diminishes downstream in the Thames estuary from 6.0 m at Tilbury to 5.7 m at Thames Haven and 5.2 m at Southend. It also diminishes northward along the Essex coast from 5.0 m at Burnham on Crouch to 4.8 m at Bradwell-on-Sea, 4.6 m at Brightlingsea, 4.1 m at Clacton, 3.8 m at Walton on the Naze and 3.6 m at Harwich.

2.  The Essex Coast The north bank of the Thames estuary is largely artificial, bordered by embankments or stone and concrete walls,

with wharves and docks at Tilbury. Downstream there is a transition to coastal features as the estuary widens, with beaches of sand, shelly gravel and shingle developing along the high tide line behind intertidal mudflats. Small estuaries, fringed by salt marshes, open to the Thames shore near Mucking and at Holehaven and Benfleet Creek, on either side of Canvey Island (>Fig. 7.17.1). On the northern side of Benfleet Creek the marshes give place to the sand and shingle beaches of Leigh-onSea. Shore banks and sea walls are extensive, and there has been much artificial raising of former marshland by the dumping of sediment dredged from the Thames estuary. Erosion is prevalent along the seaward margins of salt marshes, largely because of subsidence that leads to ­deepening of nearshore water and increasing wave energy at high tide, but partly because of boat swash. To the east a ridge of London Clay comes to the coast, and was formerly cliffed from Leigh-on-Sea through Westcliff to Southend, but the coast is bordered by sea walls, promenades and marine parades, and the former cliffs are steep bluffs behind these. Sea walls, groynes and beaches continue eastward, round Shoebury Ness, and the intertidal mudflats widen and become sandy (Maplin Sands) off the low-lying deltaic island of Foulness. The River Crouch, joined by the River Roach on the northern side of Foulness, has a wide estuary underlain by up to

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and Eastchurch Marshes. At Shell Ness, the easternmost point, shelly beach sand drifting south–east has formed a small spit. The NE coast of the Isle of Sheppey rises to bluffs in London Clay at Leysdown on Sea and a prominent cliff, also in London Clay, on Warden Point. The coast to the west consists of eroding cliffs subject to recurrent landslides and mudflows on London Clay. Sheerness, on the NW coast, bordering the mouth of the Medway estuary, has a walled waterfront. Salt marshes in the Medway estuary are undergoing dissection and shrinkage, with marginal cliffing, the tidal creeks widening and intervening areas breaking up into grassy islets. At Allhallows, a ridge of London Clay is intersected by the southern shore of the Thames estuary, where the mudflats, over a mile wide at low tide, are backed by a thin shelly sand and pebble beach and a cliff cut in ­alluvium. Upstream the Thames swings southward to Shornmead Fort, then westward to Gravesend, and the estuary shore is bordered by artificial banks of earth or

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concrete protecting low-lying reclaimed land that was formerly salt marsh or reed fen. Upstream are port and industrial installations, and the  Thames Barrier was built to exclude storm-driven high  tides from the Port of London. This Barrier has increased the risk of marine flooding downstream, leading  to the building of such structures as the Dartford Gate  at the mouth of the tributary Darent River, and surge gates on the promenades of bordering towns, as at Queenborough.

References Hutchinson JN (1968) The coastal landslides of Kent. Proc Geol Assoc 79:227–237 Lewis WV (1932) The formation of Dungeness foreland. Geogr J 80:309–324 Middlemiss FA (1983) Instability of chalk cliffs between the South Foreland and Kingsdown, Kent. Proc Geol Assoc 94:115–122

7.17  Essex 1.  Introduction Much of the Essex coast is low-lying and estuarine, and there has been extensive embanking and reclamation of former marshland. Salt marshes persist on the shores of estuaries, and on parts of the outer coast where there is a wide intertidal zone, but on the more exposed sectors there are beaches of sand, sometimes with shingle or shells. There are only a few cliffs, mainly cut in glacial drift, and some of these have been stabilised as bluffs behind sea walls, particularly at seaside resorts such as Clacton. North of the Thames estuary there are several long estuarine inlets, some of which may be inherited from former mouths of the Thames. Mean spring tide range diminishes downstream in the Thames estuary from 6.0 m at Tilbury to 5.7 m at Thames Haven and 5.2 m at Southend. It also diminishes northward along the Essex coast from 5.0 m at Burnham on Crouch to 4.8 m at Bradwell-on-Sea, 4.6 m at Brightlingsea, 4.1 m at Clacton, 3.8 m at Walton on the Naze and 3.6 m at Harwich.

2.  The Essex Coast The north bank of the Thames estuary is largely artificial, bordered by embankments or stone and concrete walls,

with wharves and docks at Tilbury. Downstream there is a transition to coastal features as the estuary widens, with beaches of sand, shelly gravel and shingle developing along the high tide line behind intertidal mudflats. Small estuaries, fringed by salt marshes, open to the Thames shore near Mucking and at Holehaven and Benfleet Creek, on either side of Canvey Island (>Fig. 7.17.1). On the northern side of Benfleet Creek the marshes give place to the sand and shingle beaches of Leigh-onSea. Shore banks and sea walls are extensive, and there has been much artificial raising of former marshland by the dumping of sediment dredged from the Thames estuary. Erosion is prevalent along the seaward margins of salt marshes, largely because of subsidence that leads to ­deepening of nearshore water and increasing wave energy at high tide, but partly because of boat swash. To the east a ridge of London Clay comes to the coast, and was formerly cliffed from Leigh-on-Sea through Westcliff to Southend, but the coast is bordered by sea walls, promenades and marine parades, and the former cliffs are steep bluffs behind these. Sea walls, groynes and beaches continue eastward, round Shoebury Ness, and the intertidal mudflats widen and become sandy (Maplin Sands) off the low-lying deltaic island of Foulness. The River Crouch, joined by the River Roach on the northern side of Foulness, has a wide estuary underlain by up to

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25 m of mud and peat deposits. The persistence of these Essex estuaries is due to continuing submergence and a low sediment yield from the rivers. From Burnham on Crouch on the north bank of the estuary an embankment round the Southminster peninsula to Bradwell on Sea and the larger Blackwater estuary. On the eastern coast of the Southminster peninsula, the salt marsh is fronted by sand and mudflats up to 3 km

wide at low spring tides. Dengie Marsh is a wide salt marsh in front of a grassy sea wall with numerous parallel creeks at right angles to the shore. The salt marsh has segments of  shelly (cockle) beach along the outer edge, which is eroded and receding (Reed 1990). In places the shelly sand has been washed in over the salt marsh as a chenier (>Fig. 7.17.2), leaving an unvegetated clay terrace on the seaward side, eroding marginally (Burd 1992).

⊡⊡ Fig. 7.17.1 Eroding salt marsh on Canvey Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.17.2 Chenier of shelly gravel near Sales Point. (Courtesy Geostudies.)

Essex

Northey Island, below Maldon in the Blackwater estuary, is a site where managed retreat has been attempted as an alternative to expensive elaboration of shore defences. The estuary of the River Colne enters from the north besides Colne Point, where a sand and shingle spit has grown NNW along the shore to Sandy Point. From Colne Point the beach-fringed coast extends eastward, St. Osyth Beach being backed by a salt marsh that has been accumulating since a basalt peat formed about 4,280 years ago (Butler et al. 1981). ⊡⊡ Fig. 7.17.3 Slumping cliff in London Clay north of Walton. (Courtesy Geostudies.)

⊡⊡ Fig. 7.17.4 Cliff cut in Pliocene Red Crag at The Naze. (Courtesy Geostudies.)

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The coast at Clacton and Frinton-on-Sea has an esplanade fronting bluffs that were formerly cliffs, bordered by a sand and shingle beach that continues past the long pier at Walton. At the northern end of the promenade the cliffs of The Naze begin. These show clayey brickearth over slumping London Clay (>Fig. 7.17.3), which is overlain by shelly sand, the Pliocene Red Crag in the cliff towards The Naze (>Fig. 7.17.4). At The Naze, a sand and shingle beach borders salt marsh islands. A shelly beach has been driven on to salt

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⊡⊡ Fig. 7.17.5 The ebb road to Horsey Island at low tide. (Courtesy Geostudies.)

marshes, the outer margins of which are eroding, and there are washovers of sand and shells deposited by storm surges on the salt marsh. Walton Channel runs north from behind Walton-on-the-Naze, and to the west is Walton Backwater and The Wade, where an ebb road (submerged at high tide) runs north to Horsey Island (>Fig. 7.17.5). Hamford Water is a submerged lowland in London Clay, a tidal embayment containing a good deal of salt marsh and soft intertidal mud, and Horsey Island is a reclaimed area surrounded by earthen sea banks. At Dovercourt the beach is retained by many groynes and the peninsula at Harwich, bordering the Stour estuary, is bordered by piers and quays. The bay Upstream there are alternations of salt marsh and bluff behind a

muddy shore exposed at low tide, then the estuary ­narrows at Manningtree. The Suffolk boundary runs along the River Stour.

References Burd F (1992) Erosion and vegetation change in the salt marshes of Essex and North Kent between 1972 and 1988. Research and Survey in Nature Conservation, 42. Peterborough, Cambridgeshire, UK Butler RJ, Greensmith JT, Wright LW (1981) Shingle spits and salt marshes in the Colne Point area of Essex. Occasional Papers in Geography, 18. Queen Mary College, London Reed DJ (1990) The impact of sea level rise on coastal salt marshes. Prog Phys Geog 14:465–481

7.18  Suffolk 1.  Introduction

2.  The Suffolk Coast

The Suffolk coast has some cliffed sectors about 10–20 m high, cut in Pliocene Crag formations overlain by glacial drift, but is otherwise generally low lying. It is interrupted by the estuaries of the Orwell, the Deben and the Alde, the latter deflected southward by the longshore shingle spit of Orford Ness. Mean spring tide ranges decrease northward along the coast from 3.3 m at Felixstowe Pier and 3.2 m at Bawdsey to 2.3 m at Orford Ness and Aldeburgh, 2.1 m at Southwold and 1.9 m at Lowestoft and Gorleston.

The large estuaries of the Stour and the Orwell are submerged valleys that meet in Harwich Harbour. The Stour is tidal upstream to Flatford Mill and the Orwell to Ipswich. The persistence of these estuaries is due to continuing land subsidence, as well as meagre sediment yield from the rivers. They are bordered by bluffs, some cliffs, salt marshes and minor beaches, with muddy slopes exposed as the tide falls. The estuary ends at Languard Point, where the coast turns northward. A shingle beach extends past Felixstowe,

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⊡⊡ Fig. 7.17.5 The ebb road to Horsey Island at low tide. (Courtesy Geostudies.)

marshes, the outer margins of which are eroding, and there are washovers of sand and shells deposited by storm surges on the salt marsh. Walton Channel runs north from behind Walton-on-the-Naze, and to the west is Walton Backwater and The Wade, where an ebb road (submerged at high tide) runs north to Horsey Island (>Fig. 7.17.5). Hamford Water is a submerged lowland in London Clay, a tidal embayment containing a good deal of salt marsh and soft intertidal mud, and Horsey Island is a reclaimed area surrounded by earthen sea banks. At Dovercourt the beach is retained by many groynes and the peninsula at Harwich, bordering the Stour estuary, is bordered by piers and quays. The bay Upstream there are alternations of salt marsh and bluff behind a

muddy shore exposed at low tide, then the estuary ­narrows at Manningtree. The Suffolk boundary runs along the River Stour.

References Burd F (1992) Erosion and vegetation change in the salt marshes of Essex and North Kent between 1972 and 1988. Research and Survey in Nature Conservation, 42. Peterborough, Cambridgeshire, UK Butler RJ, Greensmith JT, Wright LW (1981) Shingle spits and salt marshes in the Colne Point area of Essex. Occasional Papers in Geography, 18. Queen Mary College, London Reed DJ (1990) The impact of sea level rise on coastal salt marshes. Prog Phys Geog 14:465–481

7.18  Suffolk 1.  Introduction

2.  The Suffolk Coast

The Suffolk coast has some cliffed sectors about 10–20 m high, cut in Pliocene Crag formations overlain by glacial drift, but is otherwise generally low lying. It is interrupted by the estuaries of the Orwell, the Deben and the Alde, the latter deflected southward by the longshore shingle spit of Orford Ness. Mean spring tide ranges decrease northward along the coast from 3.3 m at Felixstowe Pier and 3.2 m at Bawdsey to 2.3 m at Orford Ness and Aldeburgh, 2.1 m at Southwold and 1.9 m at Lowestoft and Gorleston.

The large estuaries of the Stour and the Orwell are submerged valleys that meet in Harwich Harbour. The Stour is tidal upstream to Flatford Mill and the Orwell to Ipswich. The persistence of these estuaries is due to continuing land subsidence, as well as meagre sediment yield from the rivers. They are bordered by bluffs, some cliffs, salt marshes and minor beaches, with muddy slopes exposed as the tide falls. The estuary ends at Languard Point, where the coast turns northward. A shingle beach extends past Felixstowe,

Suffolk

where a sea wall protects bluffs cut in Pliocene Red Crag over London Clay. Numerous closely-spaced groynes retain the southward-drifting shingle beach (> Fig. 7.18.1). To the north, the coast curves into the Deben estuary. To the north of the Deben estuary at Bawdsey Manor, southward drifting beach shingle has formed a small spit. Dark brown shingle fronts the Bawdsey cliffs, cut in Red Crag over London Clay, which outcrops on the foreshore with the remains of a submerged forest. The shingle beach widens north towards Shingle Street, where shingle ridges ⊡⊡ Fig. 7.18.1 The stabilised coast at Felixstowe. (Courtesy Geostudies.)

⊡⊡ Fig. 7.18.2 The shingle beach at Slaughden, showing multiple berms. (Courtesy Geostudies.)

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are separated by swales that become lagoons at high tide. This is the mouth of the River Alde (known downstream from Orford as the River Ore), which has been deflected southward by an 18 km long barrier spit of flint shingle Orford Ness (Carr 1986). To the north this widens, and has up to 40 parallel shingle beach ridges (Green and McGregor 1990). The shingle beach (> Fig. 7.18.2) continues past Slaug­hden to Aldeburgh as a barrier that narrows to about 50 m near the northernmost of the martello towers built in the 1800s to counter the threat of Napoleonic

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i­ nvasion. On the landward side the River Ore has embankments fringed by eroding salt marshes (> Fig. 7.18.3). Albeburgh stands on a spur of high ground. A sea wall maintains the coastline, and groynes have been built to retain a protective shingle beach. There had been a long history of coastal erosion, and the town has lost streets that ran parallel to the present coastline: offshore there are remains of wells that supplied water to lost buildings. The shingle beach runs north from Aldeburgh to Thorpeness, where bluffs cut in Pliocene Crag begin. They are fronted by a

minor sand and shingle foreland, which in Roman times projected at least 200 m seaward, but has been cut away. To the north Leiston Beach is wide and sandy, with dunes in front of a low bluff, which declines to a broad valley, within which is Minsmere a former estuarine inlet, enclosed as a reedy lake behind grassy dunes and a shingle beach. The coast rises north of Minsmere to Dunwich Heath, where there are receding cliffs cut in Pliocene sandstone (Norwich Crag) and gravels (Westleton Beds), overlain by sandy glacial drift. At Dunwich the cliffs cut in Norwich ⊡⊡ Fig. 7.18.3 Eroding salt marsh beside the River Ore at Slaughden. (Courtesy Geostudies.)

⊡⊡ Fig. 7.18.4 The Dunwich cliffs showing the shingle beach bulldozed up to protect their base. (Courtesy Geostudies.)

Suffolk

Crag and glacial drift have long been receding, and much of the old town has been lost (Parker 1978). From time to time the beach is bulldozed up to protect the cliff base from storm waves (> Fig. 7.18.4). The cliffs come to an end at Dunwich, but the shingle beach continues northward in front of extensive marshes. Northward the shingle beach is capped by low dunes that widen towards the mouth of the River Blyth at Walberswick. Sand and shingle drifting southward have accumulated on ⊡⊡ Fig. 7.18.5 The mouth of Covehithe Broad. (Courtesy Geostudies.)

⊡⊡ Fig. 7.18.6 Covehithe cliffs. (Courtesy Geostudies.)

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the northern side of the harbour jetties bordering the River Blyth at Southwold to form a wide prograded beach backed by low scrubby dunes. The north cliffs rise behind Sole Bay, where they are protected by a long wall. They are up to 6 m high at Easton Bavents and receding, behind a beach of sand and gravel. Then they are interrupted by a marshy lowland with Easton Broad behind a shingle bank, and after another cliff sector is Covehithe Broad, fringed by reeds and enclosed by a sandy beach with scattered pebbles (> Fig. 7.18.5).

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Steeply sloping cliffs in Norwich Crag at Covehithe have been in rapid retreat (> Fig. 7.18.6). Benacre Broad is a lake impounded by a shingle barrier across the mouth of a small valley. To the north the cliffs revive, then pass into bluffs behind a lobate foreland (Benacre Ness), a crescentic area of sand and shingle 2 km long and up to 400 m wide, inshore from shoals and the only depositional sector on a coastline of receding cliffs (Williams 1956; Robinson 1980). North of Benacre Ness are the eroding cliffs of Kessingland, which decline behind an elongated depositional foreland. There have also been land losses at Pakefield, to the north, but sea walls have been built to halt erosion. The cliffs pass into bluffs behind sandy South Beach at Lowestoft, and Oulton Broad flows into a channel (Lake Lothing), through the centre of the town. There has been accretion of southward-drifting sand to widen the beach north of Lowestoft, where Lowestoft

Ness is the easternmost point of England. A sea wall is backed by The Dene, an area of subdued dunes, and then wooded bluffs, which pass northward into Gunton Cliffs, cut in Norwich Crag. Cliffs and bluffs continue past Corton, where the ruins of St. Peter’s church stand on the cliff top - all that is left of the village of Newton. A little to the north the Norfolk boundary crosses the cliffs.

References Carr AP (1986) The estuary of the River Ore, Suffolk: three decades of change in a long-term context. Field Studies 6:43–58 Green CP, McGregor DFM (1990) Orfordness: a geomorphological assessment. Trans Inst Br Geogr 15:48–59 Parker R (1978) Men of Dunwich. Collins, London Robinson AHW (1980) Erosion and accretion along part of the Suffolk coast of East Anglia. Mar Geol 37:133–146 Williams WW (1956) An east coast survey: some recent changes in the coast of East Anglia. Geogr J 122:317–334

7.19  Norfolk 1.  Introduction

2.  The Norfolk Coast

The coast of Norfolk is generally low-lying, with some cliffs cut mainly in Pleistocene glacial drift, apart from the Hunstanton cliffs in Cretaceous rock and some sectors of Pliocene Crag in the south-east. The Chalk outcrops in shore platforms between Cromer and Weybourne. The north-east coast is lined by dunes fronting alluvial lowlands, and the north coast west of Weybourne is an array of spits, barrier islands and salt marshes in front of bluffs marking a Late Pleistocene coastline. Much of the coast is exposed to North Sea waves, those from the north-east generating a southward drift of beach sediment on the east coast and a westward drift on the north coast, but variations in longshore drifting result from occasional northerly and north-westerly waves. South of Hunstanton, wave energy diminishes along the east and south-east coasts of The Wash, where broad sand and mud areas are exposed at low tide. Beaches on the Norfolk coast have been supplied with sand and shingle derived from eroding cliffs and from the sorting of coastal and nearshore sediment derived from glacial drift deposits (Clayton 1989). Mean spring tide range increases along the Norfolk coast from 1.9 m at Caister to 2.6 m at Winterton Ness, 4.7 m at Cromer and 6.4 m at Hunstanton before diminishing down the east coast of The Wash to 5.9 m at King’s Lynn.

Low cliffs cut in Norwich Crag extend from the Suffolk border past Hopton to Gorleston-on-Sea, beyond which they pass into stabilised bluffs behind a sea wall and esplanade, and a sandy beach with groynes along to South Pier, at the mouth of the River Yare. Great Yarmouth has been built on a large spit that deflects the River Yare southward about 5 km. The seafront has sea walls and groynes fronting a broad promenade, and the beach widens to North Beach, a minor foreland, before narrowing in front of cliffs towards Caister-on-Sea. Cliffs and bluffs continue to Newport, where they become bluffs behind the dunes of Winterton Ness. These are up to 200 m wide on a narrow foreland built by sand deposition, with a fringing sandy beach. The low-lying coast is fringed by a grassy foredune behind a sand and gravel beach from Horsey to Sea Palling (>Fig. 7.19.1). It has been repeatedly breached by storms, and is partly protected by sea walls, while groynes were built in an attempt to retain the beach. In 1991 a chain of large offshore breakwaters was inserted between Sea Palling and Eccles-on-Sea to protect the coast and induce beach accretion. The narrow coastal dune fringe continues to Cart Gap, where a sea wall has been built and then receding cliffs up to 4 m high, cut in glacial drift, rise past Happisburgh. Gradual erosion has continued, but in 1999

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Steeply sloping cliffs in Norwich Crag at Covehithe have been in rapid retreat (> Fig. 7.18.6). Benacre Broad is a lake impounded by a shingle barrier across the mouth of a small valley. To the north the cliffs revive, then pass into bluffs behind a lobate foreland (Benacre Ness), a crescentic area of sand and shingle 2 km long and up to 400 m wide, inshore from shoals and the only depositional sector on a coastline of receding cliffs (Williams 1956; Robinson 1980). North of Benacre Ness are the eroding cliffs of Kessingland, which decline behind an elongated depositional foreland. There have also been land losses at Pakefield, to the north, but sea walls have been built to halt erosion. The cliffs pass into bluffs behind sandy South Beach at Lowestoft, and Oulton Broad flows into a channel (Lake Lothing), through the centre of the town. There has been accretion of southward-drifting sand to widen the beach north of Lowestoft, where Lowestoft

Ness is the easternmost point of England. A sea wall is backed by The Dene, an area of subdued dunes, and then wooded bluffs, which pass northward into Gunton Cliffs, cut in Norwich Crag. Cliffs and bluffs continue past Corton, where the ruins of St. Peter’s church stand on the cliff top - all that is left of the village of Newton. A little to the north the Norfolk boundary crosses the cliffs.

References Carr AP (1986) The estuary of the River Ore, Suffolk: three decades of change in a long-term context. Field Studies 6:43–58 Green CP, McGregor DFM (1990) Orfordness: a geomorphological assessment. Trans Inst Br Geogr 15:48–59 Parker R (1978) Men of Dunwich. Collins, London Robinson AHW (1980) Erosion and accretion along part of the Suffolk coast of East Anglia. Mar Geol 37:133–146 Williams WW (1956) An east coast survey: some recent changes in the coast of East Anglia. Geogr J 122:317–334

7.19  Norfolk 1.  Introduction

2.  The Norfolk Coast

The coast of Norfolk is generally low-lying, with some cliffs cut mainly in Pleistocene glacial drift, apart from the Hunstanton cliffs in Cretaceous rock and some sectors of Pliocene Crag in the south-east. The Chalk outcrops in shore platforms between Cromer and Weybourne. The north-east coast is lined by dunes fronting alluvial lowlands, and the north coast west of Weybourne is an array of spits, barrier islands and salt marshes in front of bluffs marking a Late Pleistocene coastline. Much of the coast is exposed to North Sea waves, those from the north-east generating a southward drift of beach sediment on the east coast and a westward drift on the north coast, but variations in longshore drifting result from occasional northerly and north-westerly waves. South of Hunstanton, wave energy diminishes along the east and south-east coasts of The Wash, where broad sand and mud areas are exposed at low tide. Beaches on the Norfolk coast have been supplied with sand and shingle derived from eroding cliffs and from the sorting of coastal and nearshore sediment derived from glacial drift deposits (Clayton 1989). Mean spring tide range increases along the Norfolk coast from 1.9 m at Caister to 2.6 m at Winterton Ness, 4.7 m at Cromer and 6.4 m at Hunstanton before diminishing down the east coast of The Wash to 5.9 m at King’s Lynn.

Low cliffs cut in Norwich Crag extend from the Suffolk border past Hopton to Gorleston-on-Sea, beyond which they pass into stabilised bluffs behind a sea wall and esplanade, and a sandy beach with groynes along to South Pier, at the mouth of the River Yare. Great Yarmouth has been built on a large spit that deflects the River Yare southward about 5 km. The seafront has sea walls and groynes fronting a broad promenade, and the beach widens to North Beach, a minor foreland, before narrowing in front of cliffs towards Caister-on-Sea. Cliffs and bluffs continue to Newport, where they become bluffs behind the dunes of Winterton Ness. These are up to 200 m wide on a narrow foreland built by sand deposition, with a fringing sandy beach. The low-lying coast is fringed by a grassy foredune behind a sand and gravel beach from Horsey to Sea Palling (>Fig. 7.19.1). It has been repeatedly breached by storms, and is partly protected by sea walls, while groynes were built in an attempt to retain the beach. In 1991 a chain of large offshore breakwaters was inserted between Sea Palling and Eccles-on-Sea to protect the coast and induce beach accretion. The narrow coastal dune fringe continues to Cart Gap, where a sea wall has been built and then receding cliffs up to 4 m high, cut in glacial drift, rise past Happisburgh. Gradual erosion has continued, but in 1999

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⊡⊡ Fig. 7.19.1 Foredune at Horsey. (Courtesy Geostudies.)

a decision was made to allow these cliffs to evolve naturally rather than building massive sea defences. There is a sector of drift divergence near Happisburgh, where north-easterly waves generate a westward drift to­­ wards Blakeney Point and a southward drift to Great Yarmouth, but this is complicated by the frequent incidence of waves from other directions, particularly north and north-west. Cliffs cut in glacial drift extend past Bacton to Mundesley, where a sea wall protects the promenade, but recession of the bordering cliffs leaves it standing slightly forward. The cliffs continue past Trimingham, where there has been erosion from runoff and seepage as well as wave attack at the cliff base. The cliffs show classic sections for Quaternary glacial geology, including ice-rafted masses of Chalk up to 100 m long underlain by glacial drift that was displaced by the glacial readvance. At Sidestrand more displaced Chalk masses have been exposed as the cliffs retreated. At Overstrand the receding cliffs are fronted by a beach retained by a longshore wall, and on either side of Cromer the cliff sections show a sequence of Pleistocene deposits. A large tabular mass of ice-rafted Chalk, mostly horizontal but locally folded and fractured, can be seen in the cliffs to the east, capped by Weybourne Crag. At Cromer the walled sea front is bordered by a shingle beach between groynes. It includes blue and white-coated flint nodules and chalk cobbles of recent local derivation, older brown flints and a sandy foreshore over chalk, exposed at low tide. The high coastal mound west of Cromer is of gla­­ cial  drift. Cliffs cut in glacial drift overlying Chalk have been subject to long-continued recession, accompanied by

slumping and the effects of runoff and seepage. Near East Runton the cliff section is cut by thrusts (the outcome of glacier pushing), and there has been rapid recession when the cliff is cut back into the soft, seaward dipping glacial drift. At West Runton there is a wide shore platform cut in shattered Chalk, releasing flint nodules that have accumulated as a beach. The Chalk is capped by a ferruginous conglomerate, patches of which stand as micro-mesas on the shore. At the base of the cliff is the Cromer Forest Bed, a black peaty layer with shells, representing interglacial deposition and overlain by sandy and gravelly glacial drift (>Fig. 7.19.2). Towards Sheringham the cliffs rise and pass into a steep grassy artificial slope protected by a basal concrete wall that widens into an esplanade in front of the town. There is a shingle beach and Chalk outcrops on a shore platform exposed at low tide. To the west, receding cliffs in glacial drift back a shore platform in Chalk exposed at low tide (>Fig. 7.19.3). The cliffs continue to Weybourne and then swing SSW, passing laterally into bluffs behind a coastal plain up to 3 km wide that extends along the North Norfolk coast to Hunstanton. Coastal features on the north Norfolk coast include morainic and associated ice margin deposits left at the limit of the Last Glacial readvance, which ran more or less along the present coastline, transgressing locally landward (Straw 1960). These deposits have been partly rearranged in the course of Late Quaternary marine submergence and ensuing wave and tidal current action into beaches, spits, dunes, marshes

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⊡⊡ Fig. 7.19.2 Cliffs cut in glacial drift near West Runton. (Courtesy Geostudies.)

⊡⊡ Fig. 7.19.3 Cliffs in glacial drift over Chalk west of Sheringham. (Courtesy Geostudies.)

and sandy and muddy intertidal flats (Andrews and Chroston 2000). At Weybourne the shingle beach continues WNW along the seaward fringe of the coastal plain (>Fig. 7.19.4), culminating in a large compound recurved spit at Blakeney Point. Storm surges have pushed the shingle landward on to the backing salt marsh and athwart the recurved ridges that run southward, dividing the salt marsh into several

compartments. Each of these contains salt marsh (>Fig.  7.19.5). The westward growth of Blakeney Point indicates a predominance of westward longshore drifting, but eastward counter-drifting certainly occurs during phases of westerly and north-westerly wave action, and this has contributed to the shaping of successive Far Points. Dunes have formed on the western end of Blakeney Point, surmounting shingle, and are still forming on Far Point

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⊡⊡ Fig. 7.19.4 The coast west from Weybourne towards Blakeney Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.19.5 Rising tide on Blakeney salt marsh. (Courtesy Geostudies.)

where sand is winnowed from the wide sandy beach and intertidal sand areas. The older dunes are much eroded and ­disrupted by blowouts, largely because of intense rabbit grazing and burrowing. West of Blakeney Point the salt marsh fringe on the mainland coast is protected by a very wide intertidal area with inner mudflats passing seaward to sand flats, and even at high tide wave action is usually weak. Further west the salt marshes widen behind Bob Hall’s Sand. An embankment borders the channel that serves the port

of Wells-next-the-Sea, and the coast to the west consists of a large reclaimed area, formerly tidal marshland, on the Holkham Hall estate, with a seaward fringe of dunes carrying a pine plantation. The dunes are fringed ­seaward by a wide sandy area exposed at low tide, on which there have been intricate changes in the configuration of sand bars, shingle and shelly ridges and salt marshes, some sectors having gained by accretion while others have lost by erosion. There are recently-formed grassy dunes (>Fig. 7.19.6).

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⊡⊡ Fig. 7.19.6 Recently formed dune ridge near Wellsnext-the-Sea. (Courtesy Geostudies.)

⊡⊡ Fig. 7.19.7 Hunstanton cliffs, showing white Chalk over Red Chalk, with Carstone outcropping on the shore. (Courtesy Geostudies.)

The coastal dunes are interrupted by a tidal inlet Burnham Harbour - north of Overy Staithe, and to the west is Scolt Head Island, a large barrier island consisting of shingle ridges with recurves similar to Blakeney Point overlain by dunes, fronted by a sand and shingle beach that is wide at low tide and backed by salt marshes. The shingle and sand have been derived from glacial drift in the coastal area and on the sea floor.

Scolt Head Island ends in a western spit - Far Pointmuch like that on Blakeney Point, and west of this the Brancaster and Titchwell marshes lie in the shelter of shallow Brancaster Bay. Off Titchwell and Thornham areas of shingle and dunes form on the intertidal sand flats and if they survive and become stabilised, mud accretion occurs in their lee, followed by the growth of salt marsh, as on Thornham Western Island.

Lincolnshire

At Holme- next- the- Sea a sand flat that was bare of vegetation when it became partly enclosed by a shingle spit in 1858 has become a salt marsh. The prehistoric circle Seahenge was discovered here. To the SW the ground rises to the Hunstanton plateau, and on the western side are vertical cliffs cut into white Chalk over Red Chalk (>Fig. 7.19.7), then basal brown Lower Greensand Carstone behind a sandy shore with reefs of Carstone. There is a wide shore at low tide, when shoals emerge out in The Wash. To the south the cliffs decline towards Heacham and a bluff diverges landward as a low escarpment in Chalk. A shingle beach develops towards Snettisham, where it becomes a spit that has grown southward, deflecting a small river through a series of ponds. The beach then diminishes southward and the coast becomes a salt marsh as wave energy declines behind a widening intertidal area. The Great Ouse flows to The Wash north of King’s Lynn, its outlet bordered by salt marshes backed by

7.20

embankments that protect reclaimed pastureland, the ­former Terrington Marsh. There has been a long history of land reclamation along the southern coast of The Wash, where salt marshes front the most recent embankments and wide areas of muddy sand are exposed on Breast Sand. The Lincolnshire border crosses the coast 2 km east of the mouth of the River Nene.

References Andrews J, Chroston N (2000) Holocene evolution of the north Norfolk barrier coastline in the Holkham-Blakeney-Cley area. In: Lewis SG et al (eds) Norfolk and Suffolk. Quaternary Research Association, London, pp 131–147 Clayton KM (1989) Sediment input from the Norfolk cliffs, eastern England - a century of coast protection and its effect. J Coast Res 5:433–442 Straw A (1960) The limit of the Last glaciation in North Norfolk. Proc Geol Assoc 71:373–390

7.20  Lincolnshire 1.  Introduction The coast of Lincolnshire is low-lying and depositional, the Kimmeridge Clay extending beneath The Wash and the Cretaceous formations of the Yorkshire Wolds being concealed by glacial drift in the coastal region. There has been reclamation of former salt marsh and mudflats on the southern and western sides of The Wash, but the North Sea coast is sandy, backed by dunes. To the north, at Donna Nook, salt marshes appear in front of the dunes, and the sandy beach narrows behind shore mudflats in the Humber estuary. Much of the coastal sediment has been derived from drift deposits of the Last Glaciation and Holocene (Postglacial) peat and clay deposits, sorted and distributed by waves and tidal currents (Robinson 1964). Mean spring tide range at Boston is 6.8 m, at Skegness 6.1 m and at Grimsby 6.0 m.

2.  The Lincolnshire Coast The southern coast of The Wash is low-lying, with embankments fronting reclaimed land and salt marsh descending to intertidal mudflats on the seaward side. The Fenland was embanked and reclaimed in stages, particularly in the sixteenth to nineteenth centuries, when the coastline ad­vanced northward and an artificial system of drainage

canals fed the rivers (Ouse, Nene and Welland) that were confined between artificial levees. Reclamation declined during the twentieth century, but salt marsh has continued to spread seaward as accretion proceeded in The Wash (Kestner 1962). The coastline is thus artificial, and remains so, on the western shore of The Wash, north from the Witham estuary. A wide reclaimed marshland, with up to four clearly defined sea banks marking stages in historical reclamation which began in Roman times, extends from Boston to Skegness. The modern sea wall is fringed by a salt marsh, then wide mudflats out to the Boston Deeps, the broad channel that widens from the mouths of the Welland and the Witham. North of the mouth of the Witham (The Haven) is Freiston Shore. Reclaimed marshland was abandoned here in the 1990s when gaps were cut in the enclosing embankment, and meadows are changing into salt marshes and mudflats. At Butterwick Low a scour trough has developed in the gap in the embankment, and former drains leading to it have become incised, with waterfalls working headward (> Fig. 7.20.1). To the NE, Wainfleet Haven (Steeping River) flows between artificial banks, then out across the intertidal zone, which is muddy to the west and sandy to the east. The mouth of Wainfleet Haven has been deflected by the southward growth of the spit at Gibraltar Point, where the coast and foreshore become sandy as exposure

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At Holme- next- the- Sea a sand flat that was bare of vegetation when it became partly enclosed by a shingle spit in 1858 has become a salt marsh. The prehistoric circle Seahenge was discovered here. To the SW the ground rises to the Hunstanton plateau, and on the western side are vertical cliffs cut into white Chalk over Red Chalk (>Fig. 7.19.7), then basal brown Lower Greensand Carstone behind a sandy shore with reefs of Carstone. There is a wide shore at low tide, when shoals emerge out in The Wash. To the south the cliffs decline towards Heacham and a bluff diverges landward as a low escarpment in Chalk. A shingle beach develops towards Snettisham, where it becomes a spit that has grown southward, deflecting a small river through a series of ponds. The beach then diminishes southward and the coast becomes a salt marsh as wave energy declines behind a widening intertidal area. The Great Ouse flows to The Wash north of King’s Lynn, its outlet bordered by salt marshes backed by

7.20

embankments that protect reclaimed pastureland, the ­former Terrington Marsh. There has been a long history of land reclamation along the southern coast of The Wash, where salt marshes front the most recent embankments and wide areas of muddy sand are exposed on Breast Sand. The Lincolnshire border crosses the coast 2 km east of the mouth of the River Nene.

References Andrews J, Chroston N (2000) Holocene evolution of the north Norfolk barrier coastline in the Holkham-Blakeney-Cley area. In: Lewis SG et al (eds) Norfolk and Suffolk. Quaternary Research Association, London, pp 131–147 Clayton KM (1989) Sediment input from the Norfolk cliffs, eastern England - a century of coast protection and its effect. J Coast Res 5:433–442 Straw A (1960) The limit of the Last glaciation in North Norfolk. Proc Geol Assoc 71:373–390

7.20  Lincolnshire 1.  Introduction The coast of Lincolnshire is low-lying and depositional, the Kimmeridge Clay extending beneath The Wash and the Cretaceous formations of the Yorkshire Wolds being concealed by glacial drift in the coastal region. There has been reclamation of former salt marsh and mudflats on the southern and western sides of The Wash, but the North Sea coast is sandy, backed by dunes. To the north, at Donna Nook, salt marshes appear in front of the dunes, and the sandy beach narrows behind shore mudflats in the Humber estuary. Much of the coastal sediment has been derived from drift deposits of the Last Glaciation and Holocene (Postglacial) peat and clay deposits, sorted and distributed by waves and tidal currents (Robinson 1964). Mean spring tide range at Boston is 6.8 m, at Skegness 6.1 m and at Grimsby 6.0 m.

2.  The Lincolnshire Coast The southern coast of The Wash is low-lying, with embankments fronting reclaimed land and salt marsh descending to intertidal mudflats on the seaward side. The Fenland was embanked and reclaimed in stages, particularly in the sixteenth to nineteenth centuries, when the coastline ad­vanced northward and an artificial system of drainage

canals fed the rivers (Ouse, Nene and Welland) that were confined between artificial levees. Reclamation declined during the twentieth century, but salt marsh has continued to spread seaward as accretion proceeded in The Wash (Kestner 1962). The coastline is thus artificial, and remains so, on the western shore of The Wash, north from the Witham estuary. A wide reclaimed marshland, with up to four clearly defined sea banks marking stages in historical reclamation which began in Roman times, extends from Boston to Skegness. The modern sea wall is fringed by a salt marsh, then wide mudflats out to the Boston Deeps, the broad channel that widens from the mouths of the Welland and the Witham. North of the mouth of the Witham (The Haven) is Freiston Shore. Reclaimed marshland was abandoned here in the 1990s when gaps were cut in the enclosing embankment, and meadows are changing into salt marshes and mudflats. At Butterwick Low a scour trough has developed in the gap in the embankment, and former drains leading to it have become incised, with waterfalls working headward (> Fig. 7.20.1). To the NE, Wainfleet Haven (Steeping River) flows between artificial banks, then out across the intertidal zone, which is muddy to the west and sandy to the east. The mouth of Wainfleet Haven has been deflected by the southward growth of the spit at Gibraltar Point, where the coast and foreshore become sandy as exposure

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Lincolnshire

⊡⊡ Fig. 7.20.1 During the ebb a waterfall develops at the head of an incised channel along a former meadow drain at Butterwick Low. (Courtesy Geostudies.)

⊡⊡ Fig. 7.20.2 Coastal dunes at Theddlethorpe. (Courtesy Geostudies.)

to waves from the North Sea increases. Gibraltar Point has grown southward as the result of longshore drifting of sand, butit has been shaped by waves moving across a shallowoffshore zone. The Point consists of alternating layers of wave-deposited sand and shingle, and has been widened by accretion, with formation of surmounting parallel dune ridges. South from Skegness the dune ridges divide and are separated by swales, indicating spits that grew successively

in front of a wide low plain. Southward drifting of sand and shingle has supplied a series of spits that have formed and then been removed and formed again, so that the coastline has prograded eastward (King 1970). The seafront at Skegness has a promenade fronted by wide beaches of shingle and sand, but these narrow northward to Ingoldmells Point. The coast between Skegness and Mablethorpe has had a long history of erosion, and several villages and parishes have been lost. Much of the

Yorkshire

coast is now artificial, lined by walls built after the 1953 storm surge, when damage was extensive. Ingoldmells Point is a hillock of glacial drift, past which the coastline swings NNW. At Chapel St. Leonard’s, rounded mounds of eroded Boulder Clay are exposed on the shore, backed by a sandy beach and dunes. Sutton on Sea, Trusthorpe and Mablethorpe are seaside resorts with sea walls fronted by beaches of sand and some gravel and many groynes. At North End, Mablethorpe, dunes begin and widen northward to Theddlethorpe, where there is a new seaward foredune (> Fig. 7.20.2). The Saltfleetby-Theddlethorpe National Nature  Re­­ serve has a vegetation succession across parallel dune ridges, the oldest dating from the fourteenth century, with woodland and scrub on the inner part and marram grass on the younger dunes to seaward. The wide dune fringe ­continues to Saltfleet Haven behind a salt marsh that fades out along the shore, and very wide muddy sand flats.

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Patches of salt marsh go on to Donna Hook, where the coastline swings NW and there are low grassy dunes in front of a sea bank, and segments of marsh on the shore. The coast then curves west along the southern side of the Humber estuary. Cleethorpes has a yellow sand beach, and a former boulder clay cliff is now a bluff behind the ­promenade. To the west the reclaimed dock area of Grimsby begins. Lincolnshire comes to an end in the Humber estuary.

References Kestner FJT (1962) The old coastline of The Wash. Geogr J 128:457–478 King CAM (1970) Changes in the spit at Gibraltar Point, Lincolnshire, 1951–1969. East Mid Geogr 5:19–30 Robinson AHW (1964) The inshore waters, sediment supply and coastal changes of part of Lincolnshire. East Mid Geogr 3:307–321

7.21    Yorkshire

1.  Introduction The southern part of the Yorkshire coast, from the Humber estuary to Bridlington, is cut into Pleistocene glacial drift (boulder clay) deposits. The underlying Chalk rises to form the Flamborough Head promontory, and to the north the largely cliffed coastline crosses Lower Cretaceous and Jurassic formations in descending sequence. In Cleveland the sequence continues down through the Lias to the Triassic formations in the Tees valley, and then the Perm­ian (Magnesian Limestone) north from Hartlepool. Apart from the Humber estuary and the Tees valley, the coast is generally steep and cliffed, bordered by beaches and shore platforms. Mean spring tide range decreases from 6.5 m at Hull, beside the Humber estuary to 5.9 m at Spurn Head, and then northward along the coast to 5.1 m at Bridlington, 4.8 m in Filey Bay and at Scarborough, 4.6 m at Whitby and the mouth of the River Tees, and 4.3 m at Hartlepool.

2.  The Yorkshire Coast The north shore of the Humber estuary has minor sandy beaches, strips of salt marsh and embankments fronted at low tide by sloping mud with minor creeks. At Hull there are quays and docks along a walled coastline, but downstream these give place to earth walls fronting a reclaimed lowland, some sectors having salt marsh fringes.

Spurn Head is a long curving spit of sand and shingle, extending for 4.8 km. It has been supplied with sand and shingle derived from the glacial drift in the Holderness cliffs and carried south by longshore drifting, and has proba­­bly also received sediment from glacial drift deposits on the sea floor. At Kilnsea Warren there are remains of a grassy sea bank over a low cliff of Devensian boulder clay, which has been cut back by wave attack to a serrated form (> Fig. 7.21.1). Kilnsea Warren marks the beginning of the Holderness coast. Glacial drift deposits form the gently undulating low plateau behind the coast from Spurn Head to Brid­lington bordered by a cliffed coast in rapid retreat. The cliffs in soft boulder clay are steeply sloping, the drift deposits being generally too weak to sustain a vertical profile. They are receding at 1–2 m/year, and it is estimated that more than 80 square miles of land has been lost since Roman times: many villages and parishes have been lost The nearshore sea is generally muddy with clay kept in suspension by wave turbulence, the water becoming clearer seaward. Attempts have been made to halt erosion by building sea walls and dumping boulders, but it is difficult to protect 56 km of soft clay cliff from storm waves generated over the North Sea. The beaches are of sand and shingle, with some larger boulders below low tide level. The drift deposits contain varying proportions of clay, silt, sand, gravel and boulders, and typically about 30% of the material eroded from the cliffs is sand and coarser sediment of

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coast is now artificial, lined by walls built after the 1953 storm surge, when damage was extensive. Ingoldmells Point is a hillock of glacial drift, past which the coastline swings NNW. At Chapel St. Leonard’s, rounded mounds of eroded Boulder Clay are exposed on the shore, backed by a sandy beach and dunes. Sutton on Sea, Trusthorpe and Mablethorpe are seaside resorts with sea walls fronted by beaches of sand and some gravel and many groynes. At North End, Mablethorpe, dunes begin and widen northward to Theddlethorpe, where there is a new seaward foredune (> Fig. 7.20.2). The Saltfleetby-Theddlethorpe National Nature  Re­­ serve has a vegetation succession across parallel dune ridges, the oldest dating from the fourteenth century, with woodland and scrub on the inner part and marram grass on the younger dunes to seaward. The wide dune fringe ­continues to Saltfleet Haven behind a salt marsh that fades out along the shore, and very wide muddy sand flats.

7.21

Patches of salt marsh go on to Donna Hook, where the coastline swings NW and there are low grassy dunes in front of a sea bank, and segments of marsh on the shore. The coast then curves west along the southern side of the Humber estuary. Cleethorpes has a yellow sand beach, and a former boulder clay cliff is now a bluff behind the ­promenade. To the west the reclaimed dock area of Grimsby begins. Lincolnshire comes to an end in the Humber estuary.

References Kestner FJT (1962) The old coastline of The Wash. Geogr J 128:457–478 King CAM (1970) Changes in the spit at Gibraltar Point, Lincolnshire, 1951–1969. East Mid Geogr 5:19–30 Robinson AHW (1964) The inshore waters, sediment supply and coastal changes of part of Lincolnshire. East Mid Geogr 3:307–321

7.21    Yorkshire

1.  Introduction The southern part of the Yorkshire coast, from the Humber estuary to Bridlington, is cut into Pleistocene glacial drift (boulder clay) deposits. The underlying Chalk rises to form the Flamborough Head promontory, and to the north the largely cliffed coastline crosses Lower Cretaceous and Jurassic formations in descending sequence. In Cleveland the sequence continues down through the Lias to the Triassic formations in the Tees valley, and then the Perm­ian (Magnesian Limestone) north from Hartlepool. Apart from the Humber estuary and the Tees valley, the coast is generally steep and cliffed, bordered by beaches and shore platforms. Mean spring tide range decreases from 6.5 m at Hull, beside the Humber estuary to 5.9 m at Spurn Head, and then northward along the coast to 5.1 m at Bridlington, 4.8 m in Filey Bay and at Scarborough, 4.6 m at Whitby and the mouth of the River Tees, and 4.3 m at Hartlepool.

2.  The Yorkshire Coast The north shore of the Humber estuary has minor sandy beaches, strips of salt marsh and embankments fronted at low tide by sloping mud with minor creeks. At Hull there are quays and docks along a walled coastline, but downstream these give place to earth walls fronting a reclaimed lowland, some sectors having salt marsh fringes.

Spurn Head is a long curving spit of sand and shingle, extending for 4.8 km. It has been supplied with sand and shingle derived from the glacial drift in the Holderness cliffs and carried south by longshore drifting, and has proba­­bly also received sediment from glacial drift deposits on the sea floor. At Kilnsea Warren there are remains of a grassy sea bank over a low cliff of Devensian boulder clay, which has been cut back by wave attack to a serrated form (> Fig. 7.21.1). Kilnsea Warren marks the beginning of the Holderness coast. Glacial drift deposits form the gently undulating low plateau behind the coast from Spurn Head to Brid­lington bordered by a cliffed coast in rapid retreat. The cliffs in soft boulder clay are steeply sloping, the drift deposits being generally too weak to sustain a vertical profile. They are receding at 1–2 m/year, and it is estimated that more than 80 square miles of land has been lost since Roman times: many villages and parishes have been lost The nearshore sea is generally muddy with clay kept in suspension by wave turbulence, the water becoming clearer seaward. Attempts have been made to halt erosion by building sea walls and dumping boulders, but it is difficult to protect 56 km of soft clay cliff from storm waves generated over the North Sea. The beaches are of sand and shingle, with some larger boulders below low tide level. The drift deposits contain varying proportions of clay, silt, sand, gravel and boulders, and typically about 30% of the material eroded from the cliffs is sand and coarser sediment of

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the kind found on the beaches, the finer material being dispersed quickly by waves and currents (Mason and Hansom 1989). The coast curves north-eastward past Kilnsea and the boulder clay cliffs grow to between 5 m and 10 m, and locally higher. The cliff sections show pebbles and angular stones scattered through the boulder clay, and a variable capping of sandy gravel. The rate of cliff recession on this coast varies in relation to the width and height of the southward-drifting

beach, wave attack on the cliff base being impeded where the beach is wide and high, as at Sand Le Mere. These higher and wider lobes and strips of beach are separated by low sectors, termed ords by Pringle (1985), where the upper beach was missing (> Fig. 7.21.2). Hornsea Mere is a reed-fringed lake in a hollow on the undulating surface of the glacial drift where the cliffs decline across the mouth of the valley. At Hornsea cliff recession has been halted by more than a mile of promenade wall (South Cliff to North Cliff), and groynes retain ⊡⊡ Fig. 7.21.1 Scalloped cliffs in boulder clay at Kilnsea Warren, with storm surge washovers. (Courtesy Geostudies.)

⊡⊡ Fig. 7.21.2 Cliffs near Withernsea, with ords on shore. (Courtesy Geostudies.)

Yorkshire

compartments of sand and shingle in patterns produced by southward drifting. The receding boulder clay cliffs run on north past Fraisthorpe and Hilderthorpe to Bridlington, where a sea wall with a promenade fronts stabilised slopes of boul­­der clay, and the beach is of sand and shingle derived from gravel overlying the boulder clay, with a scattering of chalk cobbles and pebbles that have drifted south from Sewerby. At Sewerby an emerged beach is banked up against an old cliff. Bluffs are cut in boulder clay and glacial sand and gravel, and as the coast curves NE, Chalk rises past Danes Dyke in the cliffs of Flamborough Head. The Upper Chalk of Yorkshire is flintless, and the persistence of a promontory is related to the homogeneity of the Chalk rather than its hardness. It is much faulted, and the coastline shows many caves, arches, blowholes and inlets cut out along fault planes. The shore platform is well developed, often with a relatively steep transverse gradient (Trenhaile 1974) (> Fig. 7.21.3). The cliffs are interrupted by a cove at South Landing, where the mouth of an incised valley in Chalk, overlain by boulder clay, runs down from the village of Flambor­­ ough.  A slope-over-wall coast (> Fig.  7.21.4) continues out to Flamborough Head. The coast curves NW, and is exposed to strong wave action from the long northward fetch. On the north coast of Flamborough Head the Chalk dips gently southward, and forms an escarpment cliff with a bouldery basal ledge.  The Chalk has been dissected along major joints and faults to form scalloped cliffs with rectilinear coves and little headlands, and a variety of arches, buttresses, ⊡⊡ Fig. 7.21.3 Danes Dyke, Flamborough Head. (Courtesy Geostudies.)

7.21

stacks, ­blowholes, funnels, cauldrons, arêtes (as at Thornwick Nab), gullies, chimneys and caves. The high Chalk cliffs end north of Bempton, where the underlying Red Chalk and Speeton Clay emerge in the Speeton Cliffs as the Chalk escarpment runs inland. The coast curves NW into Filey Bay, where the Kimmeridge Clay rises into the grassy bluffs. The receding cliffs of Filey Bay are incised by small ravines between sharp-edge spurs, and fronted by a narrow shingle beach with a few boulders derived from the glacial drift, then the wide Filey Sands exposed at low tide. To the north the grassy bluffs pass into cliffs that steepen on the Lower Calcareous Grit (Middle Oolites) forming Brigg Cliff. At Filey Brigg the southwarddipping calcareous sandstone with intervening shales, has been etched into a series of reefs running out into the sea. The cliffs then run NNW, passing from the Lower Calcareous Grit to the Oxford Clay, which becomes unstable, with an apron of tumbled rock. Lebberston Cliff (>  Fig.  7.21.5) has layers of Lower Calcareous Grit over Oxford Clay, with basal Kellaways Rock. To the north, steep slumping bluffs of Oxford Clay at Osgodby Point descend to a shore platform of Inferior Oolite limestone; the Inferior Oolite Estuarine Series ­outcrops on scalloped Knipe Point, and White Nab is a  steeper headland with slumping grey cliffs of boulder clay.  Steeply sloping scrubby bluffs continue north to Scarborough. In 1993 a major landslide occurred above Black Rocks (> Fig. 7.21.6), just south Scarborough promenade, in thick boulder clay. A large amphitheatre formed, undermining Holbeck Hotel, which disintegrated and collapsed.

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⊡⊡ Fig. 7.21.4 Slope-over-wall coast at Flamborough Head. The slope is in glacial drift above a cliff cut in Chalk. (Courtesy Geostudies.)

Scarborough has a Marine Drive along the shore, sometimes awash when easterly gales drive waves over the sea wall. The cliffed headland on which the castle stands protrudes on the Lower Calcareous Grit. At Scalby Ness Rocks is the mouth of Sea Cut, a canal that runs through from the Derwent River. The coastal slope becomes higher and wider northward to Ravenscar, where a sheer cliff of sandstone rises more than 180 m above sea level at the edge of a high plateau. To the north-west is Robin Hood’s Bay, with cliffs cut in boulder clay backed by slopes rising to the Fylingdales moorland. Robin Hood’s Bay is backed by receding cliffs 50–100 m high cut in boulder clay over stratified Lower Lias limestones and shales, which outcrop in a wide intertidal shore platform with ledges of gently-dipping limestone. The hillside village of Robin Hood’s Bay has been threatened by landslides, but is protected by a high sea wall. Saltwick Nab is a bare promontory of hard shale, backed by cliffs of grey sandstone rising to Whitby Abbey. At Whitby the cliffed coast is interrupted by the mouth of the incised valley of the River Esk, and offshore the Jet Rock forms a reef extending north of the harbour. The Whitby seafront has been protected by a sea wall, and former low cliffs have been stabilised as grassy slopes. A sandy beach extends north to Sandsend, then cliffs continue behind bouldery shores and sandy beaches to the promontory at Kettleness, which stands beside Runswick Bay, cut in glacial drift that filled a Pleistocene valley. There is an upper cliff of massive Ravenscar Sandstone over steep slumping slopes in soft Lias. ⊡⊡ Fig. 7.21.5 Lebberston cliff. (Courtesy Geostudies.)

Yorkshire

7.21

⊡⊡ Fig. 7.21.6 Landslide lobe south of Scarborough. (Courtesy Geostudies.)

⊡⊡ Fig. 7.21.7 Arête on landslide south of Skinningrove. (Courtesy Geostudies.)

Staithes stands behind a cove at the mouth of a deeply incised narrow valley (locally known as a wyke) on the northern slopes of the North York Moors. To the west the coast rises to Boulby Cliff (203 m), the highest cliff on the east coast of England, capped by massive Ravenscar Sandstone over slumping Lias clays. Transverse narrow steep-sided ridges (arêtes) stand between scalloped hollows (> Fig. 7.21.7). The cliffs in Lower Lias decline between Saltburn and Redcar, and at Coatham wide intertidal sands are backed by cliffed dunes on a barrier spit that constricts the mouth of the Tees estuary. The beach and dune sands here have come mainly alongshore and from the floor of the North Sea and longshore drift, rather than from the River Tees: the sand in the estuary has been swept in from the sea by waves and rising tides. At Hartlepool a breakwater runs out, and a sea wall continues round a low promontory (The Heugh) of Magnesian Limestone. To the north low cliffs back a sandy beach along the coast north-west across the Durham border at Black Halls Rocks.

References Mason SJ, Hansom JD (1989) Cliff erosion and its contribution to a sediment budget for part of the Holderness coast. Shore Bch 56:30–38 Pringle AW (1985) Holderness coast erosion and the significance of ords. Earth Surf Proc Land 10:107–124 Trenhaile AS (1974) The geometry of shore platforms in England and Wales. Trans Inst Br Geogr 62:129–142

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Durham, Tyne and Wear

7.22  Durham, Tyne and Wear 1.  Introduction This coastline is about 32 km long, almost entirely on the outcrop of the Permian Magnesian Limestone formation. Much is cliffed, with interruptions at the mouths of incised valleys, including the lowlands of the Wear at Sunderland and the Tyne at South Shields. There are beaches of sand and shingle, which in several sectors have been augmented by waste from coal mines (the Coal measures are below the Magnesian Limestone) that has spilled or been dumped on to the shore. Mean spring tide range is 4.4 m at Seaham and Sunderland.

2.  The Coastline North-west from Hartlepool the beach is backed by dunes, which fade out as cliffs in Magnesian Limestone begin. The beach becomes intermittent between rocky outcrops, and colliery waste has drifted south-east from Crimdon Dene. At Blackhall Rocks, the coast consists of a slumping slope of boulder clay over rugged cliffs in Magnesian Limestone (>Fig. 7.22.1). The coastline is crenulate, with little coves, and cliffs that have been dissected by caves and natural arches, with an outlying tall stack. Castle Eden Dene is a deeply incised wooded valley, and to the north the cliffs are fronted by Horden Beach, which has prograded as the result of the dumping of

c­ olliery waste. Longshore drifting is mainly southward, and sorting by wave action concentrated the ‘sea coal’ that used to be collected from the shore at Horden. Foxholes Headland protrudes south of Easington Beach, which received waste from the Easington Colliery until it closed in 1992. Fine-grained sediment from the colliery waste has bound the beach sands into a terrace, the seaward edge of which has been cut back as a microcliff (>Fig. 7.22.2). There are several deep denes cut into boulder clay and locally into the underlying Magnesian Limestone. Towards Seaham there are cliffs in soft Permian sedimentary rock, overlain by glacial drift, and these are notched by several narrow incised valleys. North of Seaham Hall there are slumping cliffs in glacial drift, and a wide yellow sand and grey gravel beach, which has not been polluted by mining waste. At Ryhope the cliffs steepen on the more resistant Magnesian Limestone, fronted by structural ledges and shore platforms. Sunderland Harbour at the mouth of the River Wear is enclosed by curved jetties. Sand has been washed into the mouth of the river by waves and inflowing tides. The Permian Concretionary Limestone outcrops on the coast of Marsden Bay. The hard grey fractured limestone, with rounded or kidney shaped concretions, forms headlands and stacks, whereas the cream buff dolomite is softer and has been cut back as bays. The vertical cliffs are fronted by gravel beaches contained in successive coves, from which shore platforms with slabs and boulders ⊡⊡ Fig. 7.22.1 Slope-over-wall coast south of Blackhall Rocks, the slope in glacial drift and the cliff in Magnesian Limestone. (Courtesy Geostudies.)

Durham, Tyne and Wear

7.22

⊡⊡ Fig. 7.22.2 Easington Beach, showing an apron of colliery waste, an upper beach terrace and a microcliff. (Courtesy Geostudies.)

⊡⊡ Fig. 7.22.3 Marsden Rock, a stack and natural arch in Permian limestone, Marsden Bay. (Courtesy Geostudies.)

emerge. Lots Wife is a small stack and Marsden Rock (>Fig.  7.22.3) is larger, with many ledges of hard limestone and a natural arch fringed by boulders. The cliffbase beach is of locally-derived coarse sand and gravel. King (1953) reported that steep waves produced by onshore winds and swell were destructive here, removing beach sediment, whereas when the prevailing westerly winds blew from the land offshore they lowered and flattened the incoming swell to produce constructive waves

that generated shoreward drifting and built up the beach a the base of the cliff. There are pale grey crags of Magnesian Limestone at  Trow Quarry beach, where the cliffs have been much modified by past quarrying. The cliffs end at Trow Point, where the basal red and yellow sands of the Permian rise beneath the thick Permian limestones, followed by the underlying Carboniferous Coal Measures at the mouth of the River Tyne.

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Northumberland

At South Shields, south of the Tyne breakwaters, the beach has been widened by sand accretion, and is backed by grassy dunes. While the dunes have been built by easterly winds winnowing sand from the beach, they have also been shaped by the prevailing westerlies, blowing from land to sea. The River Tyne marks the Northumberland boundary.

Reference King CAM (1953) The relationship between wave incidence, wind direction and beach changes at Marsden Bay, Co. Durham. Trans Inst Br Geogr 19:13–23

7.23  Northumberland 1.  Introduction The Northumberland coast is dominated by Carboniferous rocks dipping generally south-east, so that the Coal Measures outcrop between Tynemouth and Amble, the underlying Millstone Grit at Warkworth and the Car­ boniferous Limestone to the north. In addition, the Whin Sill dolerite reaches the coast at Craster and Bamburgh. Much of the coast is cliffed, with interruptions at the mouths of river valleys, notably at Blyth, Amble and Berwick upon Tweed, but there is a substantial area of sand deposition north of Bamburgh and around Holy Island, where estuarine embayments are fringed by salt marsh. Mean spring tide ranges are just over 4.0 m, diminishing slightly from south to north. North Shields has 4.3 m, Blyth 4.2 m, Coquet Roads 4.3 m, Amble 4.2 m, North Sunderland 4.1 m, Holy Island 4.2 m and Berwick upon Tweed 4.1 m.

2.  The Northumberland Coast North of the mouth of the River Tyne, a sandy cove is overlooked by a ruined priory and castle on a high promontory, Gibraltar Rock, at Tynemouth. This is of Magnesian Limestone in small faulted inlier in the Coal Measures south of Whitley Bay. To the north Long Sands, backed by a sea wall extend to Cullercoats. In Whitley Bay cliffs cut in Coal Measures (sandstones with coal seams), overlain by glacial drift, have been largely walled behind Whitley Sands. An intertidal causeway leads out to St Mary’s Island, where grey Coal Measures sandstone and shales outcrop in the cliffs. At Seaton Sluice cliffs of Coal Measures sandstones fade out along the curving sandy beach to the north, which is then backed by dunes. Accretion patterns alongside groynes show that southward drifting predominates, but towards Blyth there has been some sand accretion beside the western jetty. At North Blyth the beach is backed by dunes, then

crumbling cliffs on the Coal Measure shales and ­sandstones. The cliffs decline into Newbiggin Bay, and to north the beach at Lynemouth (>Fig. 7.23.1) has been augmented by sediment from colliery waste and power ­station ash dumped on the shore and spilling from cliffed tip heaps. Druridge Bay has dunes behind a long sandy beach, with a low hinterland of boulder clay over Coal Measures. At the northern end, the High Hauxley promontory ends in cliffs cut in Coal Measures sandstones. The cliffs then decline, and dunes back the beach north-east to Amble at the mouth of the Coquet valley, where the sandstones come to an end. North of the Coquet is a hilly country on the Carboniferous Upper Limestone. The long sandy Wark­ worth Beach, between the mouths of the Coquet and the Aln, is backed by dunes, and half way along is an outcrop of Millstone Grit, seen in the wave-washed coarse sandstones in the cliff and shore boulders at Birling Carrs. Dunes have built up across the former outlet from the River Aln and beside the present mouth of the river. The sandy beach continues past Marsden Rocks, a shore platform on Millstone Grit, and cliffs and bluffs run behind Boulmer Haven. At Longhoughton are Sugar Sands, in one of several large sandy bays between boulder-strewn sandstone ledges. Howick Haven is a sandy gap in the sandstone shore platform, where sand has been washed in from the sea floor. The cliffs continue past Rumbling Kern, where there is a coastal fringe of sandstone, dipping seaward. Quarrying has left a landward-facing cliff here. On the seaward side, the sandstone cliffs are dissected by caves, arches and clefts into which waves rumble. To the north the Howick Fault crosses the shore, bringing down the Acre Limestone a kilometre south of Cullernose Point. The Carboniferous Limestone (>Fig. 7.23.2) ends abruptly against the vertical columns of the Whin Sill quartz dolerite (>Fig. 7.23.3) which disintegrates into blocks that become well rounded cannon ball stones on the shore.

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At South Shields, south of the Tyne breakwaters, the beach has been widened by sand accretion, and is backed by grassy dunes. While the dunes have been built by easterly winds winnowing sand from the beach, they have also been shaped by the prevailing westerlies, blowing from land to sea. The River Tyne marks the Northumberland boundary.

Reference King CAM (1953) The relationship between wave incidence, wind direction and beach changes at Marsden Bay, Co. Durham. Trans Inst Br Geogr 19:13–23

7.23  Northumberland 1.  Introduction The Northumberland coast is dominated by Carboniferous rocks dipping generally south-east, so that the Coal Measures outcrop between Tynemouth and Amble, the underlying Millstone Grit at Warkworth and the Car­ boniferous Limestone to the north. In addition, the Whin Sill dolerite reaches the coast at Craster and Bamburgh. Much of the coast is cliffed, with interruptions at the mouths of river valleys, notably at Blyth, Amble and Berwick upon Tweed, but there is a substantial area of sand deposition north of Bamburgh and around Holy Island, where estuarine embayments are fringed by salt marsh. Mean spring tide ranges are just over 4.0 m, diminishing slightly from south to north. North Shields has 4.3 m, Blyth 4.2 m, Coquet Roads 4.3 m, Amble 4.2 m, North Sunderland 4.1 m, Holy Island 4.2 m and Berwick upon Tweed 4.1 m.

2.  The Northumberland Coast North of the mouth of the River Tyne, a sandy cove is overlooked by a ruined priory and castle on a high promontory, Gibraltar Rock, at Tynemouth. This is of Magnesian Limestone in small faulted inlier in the Coal Measures south of Whitley Bay. To the north Long Sands, backed by a sea wall extend to Cullercoats. In Whitley Bay cliffs cut in Coal Measures (sandstones with coal seams), overlain by glacial drift, have been largely walled behind Whitley Sands. An intertidal causeway leads out to St Mary’s Island, where grey Coal Measures sandstone and shales outcrop in the cliffs. At Seaton Sluice cliffs of Coal Measures sandstones fade out along the curving sandy beach to the north, which is then backed by dunes. Accretion patterns alongside groynes show that southward drifting predominates, but towards Blyth there has been some sand accretion beside the western jetty. At North Blyth the beach is backed by dunes, then

crumbling cliffs on the Coal Measure shales and ­sandstones. The cliffs decline into Newbiggin Bay, and to north the beach at Lynemouth (>Fig. 7.23.1) has been augmented by sediment from colliery waste and power ­station ash dumped on the shore and spilling from cliffed tip heaps. Druridge Bay has dunes behind a long sandy beach, with a low hinterland of boulder clay over Coal Measures. At the northern end, the High Hauxley promontory ends in cliffs cut in Coal Measures sandstones. The cliffs then decline, and dunes back the beach north-east to Amble at the mouth of the Coquet valley, where the sandstones come to an end. North of the Coquet is a hilly country on the Carboniferous Upper Limestone. The long sandy Wark­ worth Beach, between the mouths of the Coquet and the Aln, is backed by dunes, and half way along is an outcrop of Millstone Grit, seen in the wave-washed coarse sandstones in the cliff and shore boulders at Birling Carrs. Dunes have built up across the former outlet from the River Aln and beside the present mouth of the river. The sandy beach continues past Marsden Rocks, a shore platform on Millstone Grit, and cliffs and bluffs run behind Boulmer Haven. At Longhoughton are Sugar Sands, in one of several large sandy bays between boulder-strewn sandstone ledges. Howick Haven is a sandy gap in the sandstone shore platform, where sand has been washed in from the sea floor. The cliffs continue past Rumbling Kern, where there is a coastal fringe of sandstone, dipping seaward. Quarrying has left a landward-facing cliff here. On the seaward side, the sandstone cliffs are dissected by caves, arches and clefts into which waves rumble. To the north the Howick Fault crosses the shore, bringing down the Acre Limestone a kilometre south of Cullernose Point. The Carboniferous Limestone (>Fig. 7.23.2) ends abruptly against the vertical columns of the Whin Sill quartz dolerite (>Fig. 7.23.3) which disintegrates into blocks that become well rounded cannon ball stones on the shore.

Northumberland

7.23

⊡⊡ Fig. 7.23.1 Beach with black colliery waste at Lynemouth. (Courtesy Geostudies.)

⊡⊡ Fig. 7.23.2 Slope-over-wall cliff on horizontally-bedded Carboniferous Limestone. (Courtesy Geostudies.)

Cliffs of Whin Sill quartz-dolerite extend northward past Craster to Dunstanburgh Castle. At Greymare Rock, the Sandbanks Limestone is strongly folded and there are  successive headlands of limestone and sandy bays ­northward to Seahouses. Offshore are the Farne Islands, a group of rugged dark dolerite outcrops, with columnar and deeply fissured cliffs, detached parts of the Whin Sill. They have been scraped and moulded by glaciation, and shattered by frost action.

North to Bamburgh grassy dunes occupy a wide zone behind the sandy beach, which is interrupted at low tide by the Islestone Reef, of Whin Sill quartz-dolerite. A wide beach runs past Bamburgh Castle, which is built on a thick columnar sill above grey and pink sandstones and shale. Dunes back the sandy beach, and continue along the coast to the north. In the hinterland are several north–south ridges (eskers and kames) of glacifluvial deposits.

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The sandy shore ends at Harkness Rocks, beyond which is Budle Bay, a tidal lagoon backed by gently rising glacial drift land. At low tide, a wide area of sand is exposed. The inner shore is muddy, with only a minor eroded salt marsh fringe. The floor of the bay consists of inwashed sand, threaded by tidal channels. Budle Bay is bordered northward by the Ross Peninsula, which has a  high-seaward dune fringe, Ross Links. At low tide, Lindisfarne Lagoon dries out as Holy Island Sands.

Holy Island is almost square in outline, low-lying except for dunes up to 10-m high along the northern coast (Galliers 1970). The island consists largely of Lower Carboniferous sandstones with a cover of boulder clay, and access from the mainland is by way of a sealed ebb road, which is submerged for about a third of the mean tidal cycle. The sandy northern shore curves out towards Snipe Point, where gravels emerge, then there are broad arched ledges of undulating Carboniferous Limestone at ⊡⊡ Fig. 7.23.3 Boulder beach south of cliffs in columnar Whin Sill dolerite at Cullernose Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.23.4 Whin Sill dolerite at Lindisfarne, Holy Island. (Courtesy Geostudies.)

Northumberland

Back Skerrs. Along the east coast of Holy Island, a pale grey gravelly beach stands behind broad bouldery limestone ledges, and in front of low cliffs of earthy glacisl drift, exposed in the cliff on Emmanuel Head. A dolerite intrusion, part of Whin Sill, forms an intermittent ridge along the southern coast of the island, notably at Lindisfarne Castle (>Fig. 7.23.4). North of Beal Point dunes develop behind the wide Goswick Sands. Their seaward margin is uncliffed, with marram grass spreading on to accreting sand. They include spits, bars, sand shoals, and splays. At Cocklawburn the yellow, coralline Dryburn Limestone forms Cheswick Black Rocks, its outcrop curving out from the Scremerston dunes to form bosses projecting from the intertidal sands. This part of the coast is still receiving sand washed in from sea floor shoals. Northward, the rising limestone cliffs become capped with dunes and partly fronted by grassy dunes at Cocklawburn. Limestones and shales are repeated by scroll folding on Far Skerr, Middle Sker and Near Sker, and the cliffs are fronted by projecting shore ledges of jointed pale grey limestone and sandstone, with intervening sandy coves and clefts and some cobble beaches. Towards Spittal, the strike swings northward so that the Carboniferous Limestone layers run parallel to the shore.

⊡⊡ Fig. 7.23.5 Intertidal sand spit north of Spittal, in the entrance to Berwick Harbour. (Courtesy Geostudies.)

7.23

Spittal has a little promenade, and is built on a low coastal terrace, a gravelly storm beach built when the sea stood slightly above its present level. Sand has been washed into the estuarine mouth of the Tweed by waves and inflowing tides, and a sandy cuspate foreland at Spittal culminates in an intertidal spit (>Fig. 7.23.5). Berwick is a small town beside the estuary of the River Tweed. North Pier has intercepted southward-drifting sand to form a broad accretion zone, with a wide sandy beach fronting low grassy dunes. North of Berwick the Lower Silurian rocks are fringed along the coast by seaward-dipping Carboniferous strata, as at Meadow Haven (>Fig. 7.23.6). To the north is an undulating slope-over-wall coast, the slope is grassy on glacial drift and the cliffs cut in pink and grey limestones and thin shales. The bays are cut out in softer rock, and there has been local cliff recession behind them, marked by fresh scars and slumped talus. Ribs of rocks run out across the wide rocky shore exposed at low tide, and ledges of limestone dip southward, with minor folds and intersecting strike patterns, dissected along joints. Needle’s Eye is a prominent natural arch cut through stratified sandstones. At Green’s Haven, a syncline in Carboniferous Limestone pitches north, and is seen as swirling outcrops of strata truncated on the shore. The bay

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Northumberland

⊡⊡ Fig. 7.23.6 The bay at Meadow Haven, Berwick upon Tweed. (Courtesy Geostudies.)

at Marshall Meadows is backed by high cliffs cut into shales and limestones, capped by red sandstones and conglomerates, from which some of the pebbles on the beach are derived. The cliff profile shows vertical walls on the sandstones and vegetated slopes on intervening shales. Northumberland comes to an end just north of Marshall Meadows, at the Scottish border.

Reference Galliers JA (1970) The geomorphology of Holy Island, Northumberland, University of Newcastle upon Tyne, Department of Geography Research Series, No. 6, University of Newcastle upon Tyne, Newcastle upon Tyne, p. 34

7.24 Scotland

William Ritchie · Alastair Dawson

Introduction It is only about 350 km from the northernmost part of the  Scottish mainland to the English border, and about 150 km from west to east, yet the Scottish coastline, including the numerous islands to the north and west, is about 12,000 km long, with about 15% classified as estuarine or narrow inlet, and less than 10% being some form of sandy beach (> Fig.  7.24.1). These beach and dune areas range from small coves to vast sand plains. Hard coastal forms such as plunging cliffs, pseudo-cliffs, and rock platforms form the greater part of the coastline, but there are  also good examples of such soft coastal forms as salt marshes and tidal and estuarine mudflats. The variety of the long Scottish coastline can be explained as a result of several key factors: rock type, geological structure, patterns of glaciation and deglaciation, sea level changes, and recent processes, of both the sea and the land. For these reasons an exceptionally large number of areas have some kind of conservation designation. An additional factor, of only local importance, is coastal engineering work. The variety of rock type and structure gives rise both to the general outline of the coast, where one can note a degree of parallelism that is explained by major fault lines (Steers 1973), and to particular small scale irregularities that relate to structural and lithological differences. It is  also possible to discern a broad division between the lower, sedimentary coastlines of the east, and the more rugged, steeper, fjord and island coastlines of the north and west. There are contrasts between the Pre-Cambrian sedimentary and metamorphic rocks of the north and west and the younger rocks, notably Carboniferous sedimentary rocks of the south and east. There is the imprint of Tertiary vulcanicity, with vents and a pattern of numerous sub-parallel dykes sweeping from southern Scotland NW across the peninsulas and islands of the Inner and Outer Hebrides. Major drowned estuaries and firths (Solway, Clyde, Forth, Tay, Inverness, and Dornoch) are deep indentations in the large-scale coastal planimetry, and their existence might be related to preglacial drainage systems and planation surfaces as well as fundamental geological

differences. Such broad divisions break down on the regional scale, especially where volcanic and granitic intrusions produce bold cliffs on an otherwise low land coastline. There is also the largely unresolved problem of the existence of erosion surfaces that exert a strong control of the absolute  elevation of most regional coastlines (George 1966). Glaciation and sea level changes further reduce the validity of a simple geological explanation for the physiography of the Scottish coastline. Deglaciation is recent, and the evidence of the action of ice and, more significantly, meltwater is fresh on the landscape. To some degree, the coastline is everywhere affected by the several phases of glaciation and subsequent climatic and related environmental changes. This might take many forms, but the clearest indication is derived from the striking dissimilarity between the west and east coasts. The difference is partly explained by the average position of the ice-shed, giving steep gradients and erosion-dominated, deep, fjord-like (sea lochs) coastlines to the west and north. In contrast, lower gradients for ice and meltwater movements to the east produced situations in which coastal plains with more depositional features and fewer indentations could develop. Some writers have also examined preglacial drainage patterns as a means of explaining the patterns of major inlets (Steers 1973). All around the coastline there is clear evidence of abundant deposition by ice and meltwater on the shallow sea platforms – material that a rising Holocene sea level would carry onto the lower parts of the coastline to construct great shingle strand plains, spits and bars, and extensive sand beaches. A widely held opinion is that most Scottish rivers now contribute little to the coastal sediment budget, but torrents of meltwater must have provided vast quantities of sand and gravel in the comparatively recent past. The actual extent of true cliffs, freely exposed to wave action, is less than might be expected for such a mountain and plateau-dominated country, and the contribution to the coastal sediment budget from cliff sources is largely confined to areas of sandstone and conglomerate rocks. In contrast, along many coastlines the characteristic  coastal form is the slope-over-wall profile of a thick

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 7.24.1 The coast of Scotland. 8

6

4

Palaeozoic sediment (Sandstone etc.) Precambrian gneisses etc.

Ancient metamorphics and lavas

Precambrian metamorphics and Torridon sandstone Granites 3.0

Tertiary lavas Mainly carboniferous sedimentary rocks

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Main faults Selected shallow water banks and platforms (50m contour) Maximum tidal range in metres

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Aberdeen

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Edinburgh

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Firth of Clyde

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2.8 6

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Scotland

till  cap on a steeper rock cliff, usually with a basal rock platform, and there are some examples of glacial materials forming low cliffs, especially in peripheral islands; these un­­doubtedly produce local accumulations of frequently ­coarse-grained materials. A large number of beaches, especially in the west and north are mixed lag deposits, washed out from a glacial or fluvio-glacial overburden. Salt marsh and intertidal mud and silt banks are restricted to low energy situations, especially within the firths and estuaries of the east coast and in the inner Solway Firth, but are also found as small features at the head of sea lochs and other indentations where fine sediment is available. Some evidence is available to suggest that the build-up of some estuarine mudbanks might be associated with land use changes, especially deforestation in the historic period. Another feature of special significance is the remarkable predominance of shell-derived sediment in the beach and dune areas. In the Outer Hebrides there are almost no beaches with less than 10% calcium carbonate content, and more than a quarter have shell sand contents greater than 70%. Thus some of the largest beach and dune systems in Britain are shell based, a fact that has great significance for soils, vegetation, and consequential land use (Mather and Ritchie 1977). High shell sand values are not confined to the outer islands, and there appears to be some correlation with the existence of extensive rock platform areas in the nearshore zone. Associated with shell sand dune systems, 11 sites of aeolianite (dune calcarenite) have been recorded. Every visitor to Scotland is impressed by the extent and variety of raised (emerged) coastline forms in all but the peripheral areas that lie outside the zero-isobases of the post and late-glacial raised coastlines. These peripheral areas are normally dominated by a drowned appearance, and there is some evidence from dated subtidal organic materials for postglacial submergence. Raised coastlines may consist of magnificent suites of shingle ridges 40–50 m above present sea level in exposed locations such as Jura, or they may be subtle breaks of slope in sheltered firths and estuaries. Much research has been devoted to this complex topic (Sissons 1981). Regional differences are associated with the position of the ice front during the higher relative stands of the sea as well as former coastal configurations and other geomorphological conditions. Isostatic movements also have regional patterns associated with the exact shape and position of ice sheets. Perhaps the most common form, best seen along the Firth of Clyde, the Solway, and most of the North Sea coast, is the low coastal terrace leading to one of more

7.24

degraded cliffs. Depending on lithology and altitude, these old cliffs may be gentle slopes, perhaps cut in glacial materials, or stark cliffs with caves and stacks now high and dry above an exhumed abrasion platform. The evidence of higher sea levels may be wholly or partly masked by ­subsequent deposits associated with mass movements or, more likely, by wind-blown sand. The number of such raised coastlines varies considerably, as does their clarity, but few coastlines, other than the hard rock cliff and pseudo-cliff forms, do not bear some imprint of former higher sea levels. At lower altitudes coastal deposition features formed during higher sea levels provide the basis of contemporary coastal forms whereby ridges have closed embayments as at Rattray (Walton 1956) or formed platforms upon which later landforms developed (Mather and Ritchie 1977). One of the classic areas showing sequential stages of coastal modification by depositional processes associated with a falling sea level is around the Moray Firth (Ogilvie 1923). Operating on the varied legacy of inherited coastal landforms, waves and tides give rise to a wide range of contemporary process environments. Broad distinctions can be made on the basis of exposure, which ranges from full exposure to the great fetches of the Atlantic to the moderate exposures of the North Sea coast. There are also sheltered firths an estuaries with progressive changes in energy conditions as well as extremely low energy environments in the inner sea lochs and in partially enclosed areas among the islands and peninsulas. Exposure is modified by offshore gradient; some of the most exposed coastlines (e.g. the west side of the Outer Hebrides) have extensive shallow rock platforms offshore, which reduce the violence of storm impact on the coastline. Elsewhere glaciation has produced a relatively varied sea-bed topography, but in general the west-east contrast remains valid, with shallower depths and gentler gradients occurring along the North Sea costs. One of the most important studies which was completed in 2000 for Scottish Natural Heritage, produced a comprehensive report on sediment cells around the entire coastline of Scotland. This geomorphological system is replacing formal administrative boundaries for coastal planning and management purposes. To add complexity, tide ranges vary considerably, from a minimum of about 1 m to over 7 m in the Solway. The east coast is also affected by North Sea surges, which can lead to considerable elevation of water levels and severely destructive effects as the tidal wave progresses southward down the North Sea basin. Even more variable than tides are wave conditions. Atlantic swell is important on all coastlines, even along the North Sea littoral. Most

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coastlines, however, are dominated by wind waves. Wave climatic statistics are relatively limited, but it is clear that no simple pattern of dominant wave type and direction can be described, except for coastlines that have restricted fetch or where refraction and reflection are all important. Elsewhere, the cyclonic weather patterns can produce strong winds from all directions, which set up highly variable wave inputs. Usually these are short-period waves (e.g. 5–7 seconds) and have heights of the order of 1–2 m (Buchan and Ritchie 1979), but much larger and higher wave an be experienced on most coasts during storms, several times each year. On the more open coastlines, ­offshore-onshore and longshore movements tend to be highly variable in space and time; some of the best illus­ trations of this variability are found where rivers reach the sea, and a rapidly changing geometry of spits and shallow banks is a feature of almost all Scottish river outlets. Except in bay-head situations rhythmic beach forms tend to be most common, but little research has been done on beach dynamics, and the processes are certainly compounded by variable tidal regimes and high beach water tables. Most hard rock areas are relatively inert; their characteristics arise from the position of the present interface between the sea and the former coastal slope – a slope that is more likely to have been shaped by terrestrial than marine processes. Perhaps surprisingly in a land of mountains and plateaus, true marine cliffs and active platforms are comparatively rare and are found in distinct separate regional settings. Where they exist, as on the west side of Orkney or Shetland, along the east Caithness coast, or in the spectacular columnar basalt cliffs of the Inner Heb­ rides, one can see some of the finest examples of plunging cliffs, geos, stacks, caves and arches in Europe. In summary, the long Scottish coastline is one of great physiographic and geomorphological variety. It is a complex coastline where present-day processes operate on a great diversity of inherited surfaces and landforms. Sedi­ment

supply is also variable, but there appears to be a net deficiency (although there are some notable large-scale exceptions), particularly where the inherited supply from glacial sources has been reworked onshore. Contemporary relative changes in sea level are currently a subject of considerable debate, but the meager evidence would suggest an emergence around most of the mainland and submergence of the more peripheral archipelagos. Man-made alterations in the form of groynes, sea walls, and piers are also comparatively rare except along parts of the coastline of the firths of Clyde and forth in the central lowland belt. Elsewhere such developments are small in scale and widely separated. Details of sites of geomorphological importance selected by the Joint Nature Conservation Committee are given by May and Hansom (2003). The following ten chapters deal with divisions of the coast of Scotland in an anticlockwise sequence, starting in the south-east and going round to the south-west.

References Buchan GM, Ritchie W (1979) Aberdeen Beach and Donmouth spit: an example of short term coastal dynamics. Scot Geol Mag 95:27–44 George TN (1966) Geomorphic evolution in Hebridean Scotland. Scot J Geol 2:1–34 Mather AS, Ritchie W (1977) The Beaches of the Highlands and Islands of Scotland. Countryside Commission for Scotland, Redgorton, Perth May VJ, Hansom JD (2003) Coastal Geomorphology of Great Britain. Joint Nature Conservation Committee, Peterborough, UK Ogilvie AG (1923) The physiography of the Moray Firth coast. Trans R Soc Edinb 53:377–404 Sissons JB (1981) The last Scottish ice-sheet: facts and speculative discussion. Boreas 10:1–17 Steers JA (1973) The Coastline of Scotland. Cambridge University Press, Cambridge Walton K (1956) Rattray, a study in coastal evolution. Scot Geog Mag 72:85–96

7.24.1 South-East Scotland William Ritchie 

This chapter describes the coastline from the English border at Lamberton NW past St. Abbs Head (> Fig. 7.24.1.1) North Berwick in East Lothian, then westward along the south coast of the Firth of Forth. In geological terms this runs across the eastern end of the Southern Uplands into the Midland Valley of Scotland. The Berwickshire and East Lothian coast as far as North Berwick is exposed to strong NE wave action from the North Sea, but wave energy diminishes along the southern shores of the Firth of Forth. Mean spring tide range is generally between 4 and 5 m, with a slight increase westward into the upper Firth of Forth. In Berwickshire cliffs and shore platforms continue  from the Scottish border to Burnmouth. Between Burnmouth and Eyemouth steeply-dipping sandstones outcrop in high cliffs (attaining 103 m near Fancove Head) that descend to rocky shores, and there are coves with shingle and sand beaches. North of Eyemouth the cliffs and rocky shores face NNE. St. Abbs Head is a promontory of Devonian volcanic rocks backed by a major fault that runs SE-NW. To the

north the volcanic rocks outcrop in steep, rugged cliffs up to 75 m high along a highly indented slope-over-wall coast with headlands, crags, clefts, caves, gullies, coves, reefs, stacks and skerries, formed by marine erosion dissecting along planes of weakness. This rugged coast is exposed to easterly storm waves from the North Sea (Ramsey and Broughton 2000). Silurian rocks dominate the cliffs NW to Siccar Point, site of a famous unconformity with overlying Old Red Sandstone (> Fig. 7.24.1.2). The coast declines to Pease Bay, where the tawny sand beach is backed by red cliffs. In East Lothian there are cliffs and shore platforms on limestone outcrops: at Barns Ness the limestone shore platform is dotted with rock pools, and there is a wide beach at White Sands. Dunbar has a 2 km sector of rocky coast where Old Red Sandstone outcrops in the east, basaltic tuffs in the central sector and sandstones and limestones with dykes and pipes in the west (Francis 1975). There are shore platforms at four ­levels from 25 m above to 11 m below present sea level (Gordon and Sutherland 1993): the highest platform

⊡⊡ Fig. 7.24.1.1 The rugged cliffs at St. Abbs on basaltic lavas of Middle Old Red Sandstone age.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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South-East Scotland

⊡⊡ Fig. 7.24.1.2 Siccar Point, showing Old Red Sandstone dipping seaward off Silurian rocks. (Courtesy Geostudies.)

⊡⊡ Fig. 7.24.1.3 Bass Rock, an ancient volcanic core, rises 120 m almost vertically from the sea on its exposed east face.

has been glaciated and retains a thin cover of glacial drift: it is pre-Devensian. The cliffs are bordered by stacks, and  there are coarse gravel beaches and nearshore skerries. To the west Belhaven Bay has wide wet sandy flats and a sandy marsh backed by grassy dunes on a barrier island, with recurves at the NW end. The River Tyne flows in

through reclaimed land and salt marshes on the western side. Ravensheugh Sands front cliffs in volcanic tuff and cliffs alternate with beaches along to North Berwick. Offshore is Bass Rock, a steep-sided high island, originally a volcanic plug, with a 106 m high cliff on its eastern coast (> Fig. 7.24.1.3).

South-East Scotland

At North Berwick a volcanic headland shelters the harbour and there are two sandy beaches separated by a promontory which is partly artificial. There are layered outcrops of red sandstone, tuff and agglomerate on the shore. Broad Sands occupy a bay to the west, and on Longskelly Point a basalt outcrop is backed by a raised beach 7–9 m above sea level, formed 5,000–6,000 years ago. Several small bays lie between rocky promontories on the coast west to the reef at Eyebroughty, where the coast turns SW and passes from igneous rocks at Black Point to sandstone behind Gullane Bay. South of Gullane Point is the broad, shallow Aberlady Bay at the mouth of Peffer Burn, backed by wide grassy salt marshes and some low calcareous dunes. From Aberlady Point to Craigielaw Point are cliffs cut in calciferous sandstone. Gosford Bay has a curving sand and gravel beach, backed by low grassy dunes. Sea walls dominate the coastline past the former fishing harbours of Cockenzie and Port Seton. Much of the coastline is artificial, with sea walls and port structures. Further west is the old seaside resort of Portobello, where a sea wall fronts a reclaimed area north-west to Leith Docks at the mouth of Water of Leith. The Edinburgh waterfront is largely industrial, dominated by

7. 24.1

the breakwaters of Granton Harbour and west of Granton Point low bluffs extend along the coast. Cramond Island is linked to the mainland by an intertidal causeway beside the mouth of the gravelly Almond River. The yellowbrown Drum Sands are backed by a white shelly beach. West of Hound Point the Forth Rail Bridge and Forth Road Bridge dominate the Queensferry coast. There are several piers and marinas. The land slopes down to low wooded bluffs that continue past several small coves and rocky points along to the village of Blackness, which has a gravelly and muddy shore, wide at low tide. The low bluffs come to an end at Carndon as the coastal slope passes inland behind Bo’ness, and the coastline becomes largely artificial.

References Francis EH (1975) Dunbar. In: Craig GY, Duff P (eds) The Geology of the Lothians and South East Scotland: an Excursion Guide. Edinburgh Scottish Academic Press, Edinburgh, Scotland, pp 93–106 Gordon JF, Sutherland DG (1993) (eds) Quaternary of Scotland. Geo­ logical Conservation Review Series 6, Chapman and Hall, London Ramsey DL, Broughton AH (2000) Coastal Cells in Scotland, 1 – St. Abbs Head to Fife Ness. Scottish Natural Heritage Council, Report. Battleby, p 143

479

7.24.2 Fife

William Ritchie

The Firth of Forth coast downstream from Kincardine Bridge mainly consists of low wooded bluffs that extend behind a rocky shore and intertidal sand and mudflats. In the vicinity of the Forth Railway Bridge the coast is almost entirely artificial, with sea walls. To the east is a series of rocky outcrops with low cliffs and intervening bays and inlets and shores of muddy sand, as in Dalgety Bay. The Firth of Forth then widens eastward, and a cliffed promontory faces across Mortimers Deep to the rocky island of Inchcolm. A succession of bays and promontories continues past Burntisland and the wide sandy shore of Pettycur Bay to Kinghorn, where the calciferous sandstone is interrupted by a basalt dyke. The cliffy coastline then swings north to Kirkcaldy and north-east to Leven. Largo Bay is a large south-facing, curved sandy beach backed by low amorphous dunes. The south facing coast of Fife, between Earlsferry and Fife Ness has a series of small sandy bays in gaps in the rocky shore platform. The cliffs are cut in sandstone and

limestone, with volcanic intrusions and a red conglomerate outcrops on Chapel Ness. Sauchar Point is a dark volcanic rock promontory and Cellar Dyke a harbour beside a cleft where sandstones dip seaward. East of Elie cliffs of calciferous sandstone back intermittent shore platforms, and the cliffs have a sloping profile with occasional rocky outcrops. There are shingle beaches between narrow transverse rocky reefs (>Fig. 7.24.2.1). East of Pittenweem low cliffs and shore platforms cut in back small bays. The cliffs are capped by reddish glacial drift. Most of the coastline NE of Anstruther consists of a low cliff and narrow beaches behind an intertidal wave-eroded platform on folded and faulted sedimentary structures. Crail has a small stone harbour beside a reddish sandy beach backed by cliffs. Out to sea eastward is the basaltic Isle of May, bordered by high cliffs. Fife Ness is a low undulating promontory fringed by cliffs in grey dipping limestones and shore platforms with a patchy superficial sand cover, and some dunes. The

⊡⊡ Fig. 7.24.2.1 Cliffs and rocky shore near Pittenweem.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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7.24.2    

Fife

⊡⊡ Fig. 7.24.2.2 Salt marshes fringing the Eden estuary.

⊡⊡ Fig. 7.24.2.3 Low cliffs cut into the dunes at the northern end of Tentsmuir Beach.

coastline to the NW has low cliffs and inter-tidal shore platforms with variable dip and strike. To the west are ­terraces and degraded cliffs cut into overlying boulder clay. At Boarhills volcanic rocks form the stacks known as the Rock and Spindle. St. Andrews occupies a rocky promontory with low cliffs cut in pink and grey calciferous sandstone, with an exposure of volcanic agglomerate. At low tide shore platform are exposed with strips of sand and shingle. The cliffs come to an end, and to the north is a sandy beach backed by dunes, with the world-famous links golf course.

West Sands fringe a broad spit that borders the wide estuary of the River Eden, with intertidal sand and mudflats backed by salt marshes (>Fig. 7.24.2.2). To the north Tentsmuir Sands extend for 12 km to the southern side of the Firth of Tay (Ritchie 1978). At the southern end the beach is relatively narrow and the coastline has retreated. Concrete blocks placed along the coastline as anti-tank defences in the early 1940s are now stranded out on the shore. To the north the broad intertidal sand beach of shelly sand has prograded by up to 80 m, nourished partly by longshore drift from the

Fife 

e­ roding dunes to the south and partly by the movement of sand bars derived from glacial drift deposits on the sea floor (McManus and Wal 1996). The beach is backed by a series of successively built parallel grassy dune ridges separated by marshy swales. Anti-tank blocks now stand 100 m behind grassy foredunes, indicating that progradation has continued since the 1940s (Wal and McManus 1993). Further north the beach becomes much narrower with backshore cliffing (>Fig. 7.24.2.3). Beach and dune ridges then curve round to run parallel to the southern coast of the Tay estuary, some of them truncated by backshore erosion. There are sandy shoals where the ebb and flow of tides in the Tay estuary interacts with the north-going longshore drift from the extensive beaches to the south.

7.24.2

The sandy beach comes to an end at Tayport, and the southern shores of the Firth of Tay are mainly low-lying. There are silt and sand banks and shoals to the east near the Tay Rail Bridge.

References McManus J, Wal A (1996) Sediment accumulation mechanisms on the Tentsmuir coast. In: Whittington G (ed) Fragile Environments: The  Use and Management of Tentsmuir National Nature Reserve. Edinburgh, Scotland, pp 1–15 Ritchie W (1978) The Beaches of Fife. Countryside Commission for Scotland, Perth, Australia Wal A, McManus J (1993) Wind regime and sand transport on a coastal beach-dune complex, Tentsmuir, Vol 72, Geological Society Special Publication, London, pp 159–171

483

7.24.3 Angus and Aberdeenshire

William Ritchie

On the coast between Dundee and Fraserburgh several rivers, the South Esk, North Esk, Dee, Don, Ythan and Ugie, flow into estuarine inlets or tidal basins. Glacial and fluvioglacial deposits on the shallow sea floor and sand deposition from rivers have provided sand for beaches and several large dune systems. Cliffs and shore platforms are more limited. In Angus 56% of the coastline is sandy beach, 32% rocky cliff and 12% shingle and shore platform. The north shore of the Firth of Tay borders the lowland known as the Carse of Gowrie. Barry Links occupy a triangular foreland that bears vegetated parabolic dunes with axes trending SW–NE, close to the wind resultant (Landsberg 1956). The foreland is fringed by sand and gravel beaches, converging on Buddon Ness. Low cliffs and shore platforms extend along the coast from Carnoustie, then give place to a sandy beach and dunes extending towards Arbroath. This town stands on a headland, and has a small harbour. Sandstone cliffs rise NE to The Deil’s Head, where there are several stacks.

A succession of small headlands and bays extends to Red Head. Headlands of lava stand on either side of Lunan Bay, and limestone has been quarried from the cliffs at Usan, and Elephant Rock is a stack in red sandstone. Cliffs continue to Scurdie Ness, beside the outlet from the Montrose Basin, a large tidal lagoon (3 km by 2.5 km) fed by the South Esk River. At low tide it has extensive mudflats. The coastal dunes north of Montrose extend to the mouth of North Esk River, which has an outlet deflected northward by longshore drift. North of the North Esk River grassy bluffs cut in volcanic rocks run behind a lagoon and a wide sandy plain with parallel dune ridges at St. Cyrus (> Fig. 7.24.3.1). The lagoon is part of a former channel of North Esk River that ran to a more northerly outlet that was abandoned in 1879. To the north are cliffs and stacks in basalts and andesites. At Milton Ness, the volcanic rocks give place to sandstones and the pink cliffs are fronted by rocky outcrops and pebbles on a sandy shore.

⊡⊡ Fig. 7.24.3.1 Sandy beach and dunes fronting a grassy bluff north of the North Esk River at St. Cyrus. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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7. 24. 3

Angus and Aberdeenshire

Bluffs lie behind an emerged coastal terrace, run on past the fishing village of Johnshaven, where the rocky shore is cut across narrow volcanic intrusions. The bluffs and low coastal terraces continue past a valley-mouth inlet at Inverbervie. At Crawton high vertical cliffs, with reefs protruding into the sea provide an exceptional area for viewing seabirds. Dunnottar Castle stands on a high headland bordered by a deep geo, and bluffs continue north to Stonehaven Bay. At Stonehaven red sandstones come to the coast on either side of the bay. On the northern side they run out to Garron Point, where the Highland Boundary Fault runs out across the shore, bringing up Pre-Cambrian schists and quartzites. These outcrop in cliffs northward to Findon Ness, where there is a small intrusion of granite. The micaschists return to the coast south of Cove Bay, and the coastline become more indented northwards. In Aberdeenshire the River Dee reaches the sea at Girdle Ness, and Aberdeen Harbour marks the beginning of a long sandy beach held between groynes and with a long esplanade. At the mouth of the River Don, there is a small estuary and a spit. The river mouth is unconstrained by coastal engineering works, and so migrates south and north according to the interplay of river and marine processes. Beyond the river mouth the sandy beach continues, backed by cliffed dunes. Near Balmedie the dunes are particularly mobile and are locally spilling inland. The sandy beach continues to the mouth of the Ythan estuary at Newburgh. The Sands of Forvie (>Fig. 7.24.3.2) lie north and east of the estuary, and contain an ­assemblage

of coastal dunes which have drifted from the south on to a high rocky plateau fronted by cliffs. There is a series of seven roughly parallel arcuate E–W sand ridges up to 30 m high. There is a southern area of bare sand. On the western side sand spills into the Ythan and is washed downstream. Historical records show that the drifting dunes buried Forvie chapel in the fifteenth century (Ritchie et al. 1978). North of the River Ythan the sandy beach and dunes continue until a cliffed and rocky shore capped by dunes emerges at Rockend. Andalusite schists, and glacial drift caps an ancient sloping rock platform above the cliffs at Collieston (> Fig. 7.24.3.3). At Whinnyfold a steep grassy bluff descends to a rocky shore and shingle beach, and extends round a headland to Cruden Bay. The schist outcrop ends against the Peterhead granite. In Cruden Bay an oil pipeline comes ashore to a pumping station, on its way to the refinery on the Firth of Forth. At the northern end of the bay is Port Errol, beyond which the granite cliffs become higher. The Bullers of Buchan are bold, rugged pink granite cliffs in an indented coastline exposed to easterly storm waves (> Fig. 7.24.3.4). Their outlines are the product of differential erosion along planes and zones of weakness (Buchan 1931). They include geos, blowholes, complex inlets, caves, arches, stacks, skerries and linear reefs, a range of slope-over-cliff types, and gravel and boulder beaches at the heads of inlets that end in headwalls. The granite cliffs decline to a sandy shore at Sandford Bay and large breakwaters shelter Peterhead harbour. North of the town the cliffs diminish, there are shore ⊡⊡ Fig. 7.24.3.2 The sands of Forvie. (Courtesy Department of Geography, University of Aberdeen.)

Angus and Aberdeenshire

7. 24. 3

⊡⊡ Fig. 7.24.3.3 Sloping cliffs cut in Dalradian schists at Collieston. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.3.4 Cliffs cut in granite, Bullers of Buchan. (Courtesy Geostudies.)

­ latforms on subhorizontal joint planes. Kirkton Head p and Scotstown Head are low cuspate promontories behind reefs of granite, and the granite outcrop ends against Dalradian mica schists near St. Fergus. A long sandy beach extends northward to Rattray Head, backed by high grassy dune ridges (Walton 1956).

Rattray Head is low and rocky, with boulder banks, and as the coast turns NW the dune ridges also turn in this direction behind a 5 km long sandy beach. Progradation has formed up to 18 parallel lines of dunes separated by linear swales and massive high dunes with spectacular blowouts. Behind is the shallow freshwater Loch of Strathbeg

487

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7. 24. 3

Angus and Aberdeenshire

(Bourne 1978; Ritchie et al. 1978), fringed by reedswamp and linked to the sea by a channel cut at the end of the eighteenth century. Inzie Head, to the north, has a low intertidal shore platform fronting a 10 m cliff. The beach extends past St. Combs to Inverallochy where a low bluff backs the rocky shore, with a grey cobble beach and patches of yellow sand. Fraserburgh Bay, facing NNE, has a yellow sandy beach backed by slightly undercut dunes and on the western side the seaport of Fraserburgh stands behind a broad promontory of schist, running out to Kinnairds Head.

References Bourne WRP (1978) The Loch of Strathbeg. Nature 242:93–95 Buchan S (1931) On dykes in east Aberdeenshire. Trans Edinburgh Geol Soc 12:323–328 Landsberg SY (1956) The orientation of dunes in Britain and Denmark in relation to winds. Geogr J 122:176–189 Ritchie W, Rose N, Smith JS (1978) The Beaches of Northeast Scotland. Department of Geography, University of Aberdeen, Aberdeen, Scotland Walton K (1956) Rattray: a study of coastal evolution. Scot Geogr Mag 72:85–96

7.24.4 Fraserburgh to Moray Firth and Duncansby Head

William Ritchie

The north-facing coast, west to Inverness, is exposed to N and NE wave action, the latter generating a westward longshore drift of beach sand and gravel. The eastern part is mainly cliffed, and west of Spey Bay mainly depositional. Mean spring tide range increases gradually westward from 3.1 m at Banff to 3.2 m at Whitehills 3.4 m at Buckie 3.5 m at Lossiemouth and Burghead and 4.2 m at Inverness. The coastline truncates geological formations of the Highlands, which strike SSW-NNE, so that strikingly different outcrops are crossed in rapid sequence. The coast west from Fraserburgh consists mainly of bluffs behind shore platforms. Grassy bluffs continue past Quarry Head to Aberdour Bay (>Fig. 7.24.4.1). The bluffs become cliffs on headlands, as on Strahangles Point, and there are sectors of abrasion shore platform. Sandstone cliffs and bluffs occur in Pennan Bay, which has a beach of sand and shingle. West of Pennan is Cully­ khan Bay, facing east and bordered by shore platforms.

This gives place to slates and schists on the bold cliff at Troup Head, rising 112 m above the sea, and round Crovie Head into Gamrie Bay. At Gardenstown prominent cliffs extend past More Head and westward past Head of Garness. Towards Macduff the cliffs are dissected into coves, headlands and stacks, and fronted by abrasion platforms. Macduff and Banff are towns on either side of the mouth of Deveron River. Cliffs continue westwards behind Boyndie Bay, and a sand and shingle beach extends past the mouth of the Burn of Boyndie. A low terrace fronts grassy bluffs that run out to cliffy Knock Head, a landform association, resulting from emergence, that is typical of the coast west from Banff to Portgordon. Cliffs continue westward past Whitehills across the incised valley of the Burn of Boyne, and on round to Portsoy. Sandend, beside the mouth of the Burn of Fordyce, has a wide sandy beach, backed by grassy dunes. Findlater

⊡⊡ Fig. 7.24.4.1 Boat Shore Cove, Aberdour Bay. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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7.24.4

Fraserburgh to Moray Firth and Duncansby Head

Castle, built in 1655, stands on a knoll in front of the cliffs, and to the west is a double bay at Sunnyside, with a sandy beach and grassy dunes in front of two small valleys. On the western side cliffs run out to the Dal­ radian quartzite promontory at Logie Head. At On the eastern shore of Cullen Bay grassy bluffs are fronted by a low terrace and craggy stacks of sandstone, The Three Kings and Boars Craig. A sandy beach curves out to Scars Ness. Portknockie stands behind steep cliffs and grassy bluffs above a rocky shore on Old Red Sandstone that dips steeply southward. There are several stacks, one of which is the Bow and Fiddle Rock, a natural arch (> Fig. 7.24.4.2). From Portknockie the cliffs run WSW, and declining to  steep bluffs behind a low emerged coastal terrace. Findochty and Portessie stand on this terrace. Sandstones reappear at Buckie, and to the west the grassy bluff and low coastal terrace run on past Buck­ pool, where a grey shingle beach curves out to Portgordon harbour. Another long shingle beach runs WNW to the mouth of the River Spey, as the grassy bluff recedes inland behind a widening coastal plain. Shingle beaches and barriers around the mouth of the River Spey are unusual in that they are still receiving fluvial gravel from this fast-flowing river (Gemmell et al. 2001). They form a seaward fringe to a broad emerged coastal plain with beach ridges built by northerly storm

waves on this prograded coast. The mouth of the River Spey has been deflected eastward by longshore drifting of shingle (Grove 1955). West of the Spey mouth a shingle beach fronts Lossie Forest. Between Lossiemouth and Burghead the coast is cut into desert sandstone. Burghead also stands on a narrow sandstone promontory, and a sandy beach backed by dunes extends behind Burghead Bay to the west, curving round to Findhorn. West of Findhorn Bay is a broad sandy lowland occupied by Culbin Forest. The coast at Culbin is an emerged sand and gravel plain including many spits that grew westward, backed by a prominent emerged cliff. The gravelly ridges are overlain by an extensive, formerly mobile, dune area, with dunes up to 30 m high (Comber et al. 1994). The sand is quartzose, with some feldspars, and has come largely from glacial drift. A wide mainly sandy intertidal zone with bars and spits extends from Findhorn westward to Nairn. The eastern part of the Culbin foreland is eroding, with cliffed dunes near the mouth of the River Findhorn and the western part is accreting by westward growth. At Nairn the coastline curves NW, the beach becoming a longshore spit out to Whiteness Head. This is a prograding shingle spit about 3.5 km long, ending westward in recurved ridges into salt marsh and intertidal sand flats (Ogilvie 1923; Bentley 1995; Stapleton and Pethick 1996).

⊡⊡ Fig. 7.24.4.2 Bow and Fiddle Rock at Portknockie. (Courtesy Geostudies.)

Fraserburgh to Moray Firth and Duncansby Head

To the west a sandy bay curves round to Fort George, on a promontory of Old Red Sandstone where Moray Firth narrows westward to the mouth of the River Beauly (Ogilvie 1914). Beauly Firth is a shallow estuarine basin with extensive sand and mudflats with scattered boulders exposed at low tide. From Inverness to Duncansby Head the coast faces generally SE, and is exposed to wave energy that increases northward as the SE fetch lengthens. Cromarty Firth and Dornoch Firth are relatively sheltered inlets. Mean spring tide range is 3.6 m at Cromarty and 3.8 m at Invergordon, then diminishes northward to 3.5 m at Portmahomack, 3.4 m at Golspie and 2.9 m at Wick. The low-lying northern shore of Beauly Firth is bordered by gravelly beaches and extensive mudflats exposed at low tide. The shingle beach continues past Kilmuir, where there is a cuspate foreland. At Avoch the beach becomes sandy. The low triangular protrusion of Chan­ onrie Point has shingle on the western side, while on the eastern side dunes rise above the shingle. Cliffs rise eastward from Rosemarkie along the southern coast of the Black Isle. Steep slopes descend to a low terrace with basal cliffs, a gravel beach and rocky boulderstrewn shore. The steep slopes of the Sutors of Cromarty (South Sutor) mark the beginning of Cromarty Firth. The cliffs decline to the low-lying southern shore of Cromarty Firth, where an emerged bluff is promi­ nent behind a coastal terrace, fronted by shingle. East of

⊡⊡ Fig. 7.24.4.3 The east coast of Tarbat Ness. (Courtesy Geostudies.)

7.24.4

Invergordon is wide Nigg Bay, with extensive sand flats exposed at low tide. From North Sutor the cliffs pass into steep bluffs behind a narrow emerged coastal terrace extending NNE on an almost straight coastline that follows the trend of the Great Glen Fault. Scrubby bluffs run behind an emerged bay north of Ballintore, with dunes in the Hilton of Cadboll, then continue past the village of Rockfield and on to Tarbat Ness, a long, low promontory of sandstone (> Fig. 7.24.4.3) on the southern side of the entrance to Dornoch Firth. Dornoch Firth has much more sand than Cromarty Firth. To the SW a sandy shore widens at low tide into Inver Bay, where there are two dune-capped barrier islands bordering Morrich More. These include machair, freemoving yellow parabolic dunes and stabilised grey dunes on a low-lying sandy foreland extending west to Tain (Hansom and Leafe 1990). West of Tain on the southern shore of Dornoch Firth a narrow projection, the Ness of Portnaculter, runs out across the intertidal sand flats. This is not a spit, but a gravelly esker. On the northern shore of Dornoch Firth a sandy shore fronts steep hills. The slopes decline to the east and Dornoch Point is a recurved spit that marks the northern side of the entrance to Dornoch Firth. North of Dornoch the coast is low-lying, with sandy beaches and segments of shore platform. At Embo the sandy beach is backed by dunes and to the north the beach

491

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7.24.4

Fraserburgh to Moray Firth and Duncansby Head

⊡⊡ Fig. 7.24.4.4 Loch Fleet. (Courtesy Geostudies.)

becomes a spit bordering Loch Fleet, with traces of former  recurves marking its northward growth. Loch Fleet (> Fig. 7.24.4.4) is a small lagoon that dries out to sandy flats. Further north Golspie stands on a narrow coastal plain backed by mountains. A narrow lowland fringe occurs along the coast from Brora to Helmsdale but the underlying geology is mainly conglomerates and sandstones. The narrow coastal plain fades out at Helmsdale, and the cliffs rise to the 230 m high Ord of Caithness. There are caves and coves cut out along joints in the cliffs that continue past Sarclet Head to Helman Head and Wick, in a small bay at the mouth of the Wick River. They resume northward to Noss Head, which is dissected along vertical joints into rectilinear headlands and inlets. The coastline then turns sharply westward on the southern side of Sinclair’s Bay, which a sandy surf beach (nearshore waves are up to 6 m high) backed by grassy dunes and coastal slopes. The shore becomes rocky near Keiss, where slope-over-wall escarpment cliffs are fronted by shore platforms where the seaward dip produces many parallel scarplets. The cliffs become higher on Tang Head and Freswick Bay is inset at the mouth of Gill Burn, but from Skirza Head to Duncansby Head is a 6 km sector of flat-topped cliffs, some with upper convex slopes. They show features related to the distinctive flagstone strata (Crampton and Carruthers 1914). This coast is exposed to strong wave action. Skirza Head has vertical cliffs cut in stratified yellow and red sand­ stones descending to inclined abrasional shore platforms

and trenched by many geos, notably the deep Wife Geo, a 250 m inlet with vertical walls, buttresses, pinnacles, arches and caves. Further north are higher cliffs, steep to vertical bordered by shore platforms up to 100 m wide and the three sharp-pointed Stacks of Duncansby. Just south of Duncansby Head is the elongated Duncansby Geo. Pentland Firth to the north has strong (up to 12 knot) tidal currents, and there are often stormy seas.

References Bentley M (1995) Whiteness Head. Scottish Natural Heritage, Perth, Australia Comber DPM et al (1994) Culbin Sands, Culbin Forest and Findhorn Bay. Scottish Natural Heritage Research, Survey and Monitoring 14, Edinburgh, Scotland Crampton CB, Carruthers RG (1914) The Geology of Caithness. Memoir of the Geological Survey of Great Britain, London Gemmell SGL, Hansom JD, Hoey T (2001) River-coast sediment exchanges: the Spey Bay sediment budget and management implications. In: Packham JR et al (eds) ecology and Geomorphology of coastal shingle. Westbury Academic and Scientific Publishing, Otley, West Yorkshire, pp 159–167 Grove AT (1955) The mouth of the Spey. Scot Geogr Mag 71:104–107 Hansom JD, Leafe RN (1990) The geomorphology of Morrich More. Nature Conservancy Council, Report 1161 Ogilvie AG (1914) The physical geography of the entrance to Inverness Firth. Scot Geogr Mag 30:21–35 Ogilvie AG (1923) The physiography of the Moray Firth coast. Trans R Soc Edinburgh 53:377–404 Stapleton C, Pethick J (1996) Coastal processes and management of Scottish estuaries, I: The Dornoch, Cromarty and Beauly/Inverness Firths. Scottish Natural Heritage Review 50, Edinburgh, Scotland

7.24.5 T  he Orkney and Shetland Islands

Alastair Dawson

The Orkney Islands (more than 70 of them) are largely of Middle Old Red Sandstone, with extensive cliffed coasts. There are sand and shingle beaches in some bays,  and a variety of spits, tombolos, and lagoons. In the Orkneys ayres are shingle barriers built across the mouths of small bays and oyces are the lagoons thus enclosed. To the NNE, beyond Fair Isle, are the similar Shetland Islands. The surrounding waters are often stormy, and high wave energy occurs frequently, particularly on the western (Atlantic) coasts. Mean spring tide range is generally between 2.0 and 3.0 m.

1.  Orkney Islands South Ronaldsay consists of Devonian sandstones and  flagstones, with many dolerite dykes. There is granite  on the south-eastern shore of Burray. To the NW, almost enclosed by islands is Scapa Flow, there is a sea area 24 km by 11 km with strong tidal currents. It is bordered by ­generally low coasts. In 1940, the four Churchill Causeways were built to close the North Sea entrance. Flotta is a cliff-edged relatively flat island of Middle  Old  Red Sandstone, rising to 58 m above sea level.  On its south coast, the cliffs are up to 30 m high. Fara  Island, which has 10–20 m cliffs, and the indented eastern coast of the island of Hoy are also in Middle Old  Red Sandstone, which extends round much of Scapa Flow. Hoy has high cliffs on horizontally bedded Old Red Sandstone, with small outcrops of Devonian basalt on its west coast, a coast exposed to high-energy Atlantic swell and some of the highest storm wave energy in Britain. It shows various types of cliff, caves, arches, geos, stacks, shore platforms, and cliff-top wind and spray scouring.

Horizontal bedding planes, joints, and faults have guided dissection and influenced the shape of cliff features, including the 137 m stack known as the Old Man of Hoy (> Fig. 7.24.5.1) (Hansom and Evans 1995). The SW coast also has high cliffs with some waterfalls, declining to sandy Rack Wick Bay. Mainland of Orkney is a mass of Old Red Sandstone with a small intrusion of foliated granite at Stromness and  many dykes of dolerite. There are spectacular cliffs between 20 and 60 m high, locally attaining 100 m (> Fig. 7.24.5.2), with caves, arches, geos, and stacks. As on the west coast of Hoy, there has been cliff-top wind

⊡⊡ Fig. 7.24.5.1 The Old Man of Hoy is a 100 m high, red sandstone stack on the high energy, exposed to Atlantic coast of Hoy. (Courtesy Sigurd Towrie.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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7.24.5

The Orkney and Shetland Islands

⊡⊡ Fig. 7.24.5.2 Huip Bay on Stronsay is a typical sheltered bay on the low islands of north Orkney. It is shallow, with a fringe of low dunes. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.5.3 Noup Head on Westray faces NW into the Atlantic Ocean. The cliffs, in horizontal sandstone, are typical of the west coast of most of the Orkney archipelago. (Courtesy Department of Geography, University of Aberdeen.)

and spray scouring, while horizontal bedding planes, joints, and faults have guided dissection and influenced the shape of cliffs. The west coast is almost straight. Skara Brae, behind Bay of Skaill, is a cove with a sandy beach backed by dunes, which overran an Iron Age settlement 4,500 years ago. It was excavated in 1850. To the north, a dolerite dyke forms the peninsula of Brough Head. The other coasts of Mainland are much embayed, with irregular promontories. Cliffs are lower on the more sheltered side, and are locally in cut in peat and boulder clay. There are a few sandy bays and Kirkwall is at the head of an inlet

opening into Wide Firth. To the east are bays backed by sandy beaches, and in the SE Point of Ayre, a strip of Devonian lava outcrops. Shapnisay, to the east of Mainland, has shingle barriers and a spit (Ling Holm) enclosing the tidal Ouse and a lake, Lairo Water. The island consists mainly of Old Red Sandstone, with some volcanic rocks. Much of the coastline is low with a fringe of dunes on boulder clay cliffs. Rousay has a southern fringe of dolerite at Brinyan and prominent dykes to the north, toward Sacquoy Head. The island has high bordering cliffs, especially on the west

The Orkney and Shetland Islands

coast, with geos and natural arches and a wide rocky shore platform at Quoynalonga Ness. Stronsay and Eday have irregular low cliffs cut in Old  Red Sandstone, with several curved sandy bays (> Fig. 7.24.5.2). Some smaller islands are tied by shingle barriers (ayres), as at Papa Stronsay. Westray high cliffs on its western coast, north to Noup Head (> Fig.  7.24.5.3). On the south coast, the Bay of Tuquoy has sands a kilometre wide exposed at low tide. Papa Westray is smaller but similar. Sanday, to the east, is also cliffed on its western coast, but has many sandy beaches in bays and inlets of other aspect. Central Sanday is generally low-lying, with rocky shores out to Tres Ness and easternmost Start Point, now cut off by the sea at high tide (Ayre Sound), but formerly linked by an isthmus. Quoy Ayre attaches Els Ness to the mainland, separating Little Sea, an intertidal lagoon (Mather et al. 1974). Beaches and dunes of shelly sand overlie gravels, the dunes grading landward into machair. Beaches face in various directions, some behind bays with broad intertidal zones. North Ronaldsay is a flat windswept island with low cliffs and wide rocky shores.

2.  Shetland Islands The Shetland Islands are geologically more varied than the Orkneys. Sumburgh Head is on a southern peninsula of Old Red Sandstone, with an intrusion of foliated granite to the west, outcropping on the Bay of Quendale. This

7.24.5

gives place to Dalradian slates and schists on the southern ­peninsula and on the coast past Fitful Head, where long steep slopes descend to low cliffs. St. Ninian’s Tombolo is the largest sandy tombolo in Britain, 500 m long, linking the small St. Ninian’s Isle to the SW coast of Shetland Mainland. It separates Bigton Wick, the bay to the north, from St. Ninian’s Bay to the south. Its symmetrical curved shorelines have been shaped by refracted waves arriving from north and south, and the central part is submerged by the highest spring tides. It consists of calcareous sand, piled into grassy dunes at each end, and the sand is underlain by shingle (Flinn 1974). North of Maywick, the steep west coast (Clift Hills) runs NNE, facing parallel sounds and East and West Burra Island, elongated in the same direction. Dalradian slates and schist extend along Clift Sound and out on West Burra. The southern tip of West Burra is on an intrusion of foliated granite. North of Scalloway, there are narrow, parallel SSW-NNE outcrops of Dalradian schist and an intrusion of foliated granite threading limestone, and forming long narrow sounds and peninsulas. Granite with some narrow dominates the peninsula south to Skelda Ness, but the entrance to Gruting Voe marks the start of Old Red Sandstone, which outcrops extensively round the shores of the west Zetland peninsula and Vaila island. Papa Stour is a small, relatively low island off the west coast of Shetland Mainland, separated by the Sound of Papa. It consists largely of Middle Devonian igneous rocks. The northern and eastern coasts are deeply embayed, with long inlets, known as voes. Along the cliffed western coast (> Fig. 7.24.5.4), shore platforms are rare, the cliffs ⊡⊡ Fig. 7.24.5.4 Dissected cliffs and a natural arch on Papa Stour. (Courtesy Department of Geography, University of Aberdeen.)

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7.24.5

The Orkney and Shetland Islands

plunging into deep water close inshore (Flinn 1964). There are storm-piled gravel and boulder beaches within bays. Muckle Roe is a mainly granite island with a small outcrop of Dalradian schist on the east coast and diorite bordering the narrow strait to the north-east. To the north at Brae is the narrow isthmus of Mavis Grind, between the Atlantic to the west and Sullum Voe to the east. In North-western Shetland Mainland, cliffed promontories run out west to Esha Ness, where rugged sandstone cliffs and stacks front a headland 62 m high with lava boulders. The Villans of Hamnavoe is a 3 km cliff up to 45 m high cut in Devonian volcanic rocks where storm waves have thrown boulders up on to the cliffs and cut scour features up to 15 m above sea level and up to 100 m inland (Hansom 2001). These are not regarded as tsunami features, although there was a tsunami about 7,000 years ago. To the north, Ronas Voe is a deep, steep-sided fiord (firth). There are many stacks and natural arches. On the east coast, slope-over-wall cliffs face Yell Sound, and rocky shores run round headlands and into firths. There is a rare salt marsh at Voxter. The Ayres of Swinister are shingle barrier spits and tombolos between the island of Fora Ness and a peninsula on the north-east coast of Mainland Shetland. Between the two ayres is a tidal basin, The Houb, containing drowned peat deposits indicating that this is a submerging coast. There are also remnants of salt marshes. Dalradian slates and schists dominate peninsulas and loch shores to the east, running out on the Lunna Ness peninsula, then south to Dury Voe, a wide inlet. They continue round into South Nesting Bay. South of Dales Voe, a fault brings up Old Red Sand­ stone, which forms coastal outcrops south past Lerwick and on Bressay and the Isle of Noss to the east. Sandstone cliffs with ledges skirt 180 m high Noss Head. South of Lerwick is a narrow peninsula of Old Red Sandstone, and there is a narrow Dalradian schist ourcrop at Fladdabister. The island of Mousa to the east is also of Old Red Sandstone. The island of Yell consists of Moine gneiss, and is surrounded by cliffs except for a beach at Sandwick on the west coast. Much of the island is covered by blanket bog. Unst Island has rocky scree shores and a high cliff on Herma Ness at the northern end. Outlying Muckle Flugga, an island with a lighthouse, is often said to be the northernmost part of Britain, but in fact there is another very small islet, Out Stack, a little to the north. Balta Island, a small island off the east coast of Unst, consists mainly of gabbro, and has a low west coast sheltered by Unst and a higher east coast with cliffs rising to 45 m and dissected by geos and arches. Grassy calcareous dunes

extend across the island from a NW beach through a low col, thinning toward the eastern cliffs (Mather and Smith 1974). The machair sand sheet has been dissected by rill ­erosion and deflation following overgrazing by Zrabbits. Erosion has exposed underlying dune calcarenite, an unusual formation in this region. Fetlar is a cliffy island to the south. The 13 sq. km island of Foula lies in the Atlantic Ocean 22.5 km west of Shetland Mainland, and rises to The Sneug (418 m). Its cliffy coastline is cut mainly in Old Red Sandstone (Blackbourn 1985). The west and south-west coasts, exposed to the strongest and most frequent Atlantic swells and storm waves, have high sheer cliffs, attaining 376 m on The Kame (> Fig. 7.24.5.5). On the east coast, the cliffs are much lower (20–30 m) along a fringe of faulted igneous and metamorphic  Dal­ radian rocks, forming an irregular coast with small geos ⊡⊡ Fig. 7.24.5.5 The Kame of Foula is one of the most prominent cliffed headlands in Britain, cut in steeply dipping Old Red Sandstone. (Courtesy Department of Geography, University of Aberdeen.)

The Orkney and Shetland Islands

excavated along joints. Beaches are rare, the largest being the Hiora Wick boulder beach in a bay on the north coast. Thirty-nine kilometres SW of Sumbrugh Head in southernmost Shetland is Fair Isle, of Lower Old Red Sandstone. It has an intricate, cliffed coastline. On the west coast, the cliffs rise to 150 m and on the east coast Bu Ness is attached to the island by a depositional isthmus separating North and South Haven.

References Blackbourn GA (1985) Geological Field Guide to Foula. Britoil corporate exploration stratigraphic laboratory, shetland

7.24.5

Flinn D (1964) Coastal and submarine features around the Shetland Islands. Proc Geol Assoc 75:321–340 Flinn D (1974) The coastline in Shetland. In: Goodier R (ed). The natural environment of Shetland, Nature Conservancy Council, Edinburgh, Scotland, pp 13–25 Hansom JD (2001) Coastal sensitivity to environmental change: a view from the beach. Catena 42:291–305 Hansom JD, Evans DJA (1995) Scottish landform examples – 13. Old Man of Hoy. Scot Geogr Mag 111:172–174 Mather AS, Smith JS (1974) Beaches of Shetland. Department of Geography, University of Aberdeen, Aberdeen Mather AS, Smith JS, Ritchie W (1974) The beaches of Orkney. Department of Geography, University of Aberdeen, Aberdeen

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7.24.6 North and North West Scotland (Duncansby Head to Cape Wrath and Fort William)



Alastair Dawson

1.  Introduction The north-facing coast of Scotland is generally cliffed, with some sandy bays. It is exposed to a strong wave action from north and north–west, with heavy swell originating from Arctic Ocean storms. In the eastern sector between Duncansby Head and Dunnet Head, the Orkney Islands afford some protection from northerly wave action. Mean spring tide range is generally about 4.0 m (Scrabster 4.2 m, Rispond in Loch Eriboll 3.9 m and 4.0 m at the entrance to the Kyle of Durness).

2.  The North Coast The Caithness coast extends west from Duncansby Head past John o’Groats. The Old Red Sandstone cliffs are cut by chasms and arches, and cliff falls result in slow recession. They extend along the coast from Dunnet Bay to Thurso Bay and on past Sandside Bay to the eastern side of Melvich Bay, and consist of horizontally bedded flagstones, quarried as thin, flat slabs used for paving stones. On the Dunnet Head promontory, the cliffs are up to 90 m high. They decline into Dunnet Bay, and a shingle beach gives place to calcareous sand at the head of the bay (Ritchie and Mather 1970). The southern shore is rocky and gravelly, with abandoned flagstone Quarries. Cliffs are widespread with shore platforms cut across strata that dip in various directions. This coastline then diminishes to low bluffs and curves round to Thurso. The coastline turns westward, and sandstone cliffs continue to Brims Nest with widening shore platforms. Dounreay stands behind a shelving rocky shore, and to the west a small bay, Sandside Bay, is backed by dunes that have drifted inland to bury Reay church (Ritchie and Mather 1969). On the Sutherland coast, Melvich Bay is backed by grassy dunes. Portskerra to the west has an intricate rocky

shore fronting grassy bluffs. There is a sector of Old Red Sandstone at Baligill, then Strathy Bay is cut into another granite outcrop. To the west, cliffs in Moine schist and granite run out to Strathy Point (>Fig.  7.24.6.1), which has caves and a natural arch. The schists outcrop in irregular cliffs that run south–west from Strathy Point into Armadale Bay, where the rocky shores are interrupted by a sandy beach, backed by dunes (> Fig. 7.24.6.2). Bettyhill stands on a ridge overlooking the estuary of the River Naver. The valley has three glacifluvial terraces (Ritchie and Mather 1969). Torrisdale Bay to the west has a sandy beach backed by dunes, with spits protruding at each end. Beach and dune sands are derived from glacifluvial sand and gravels. Dunes extend over the NW slopes of Drum Chuibhe. Cliffs cut in schist and gneiss extends west from Torrisdale Bay, and at Skerray the coast crosses an outcrop of serpentine. Offshore, Neave Island and Seal Island are high rocky islands of sandstone. Cliffs in Moine schist curve into Tongue Bay. The Kyle of Tongue is a long firth with cliffs that pass southward into steep bordering slopes cut in schist. At low tide, extensive sand and gravel shoals are exposed, and a low tide sand bar extends to the Rabbit Islands. To the west at Talmine, there are white sandy beaches, and the schistose cliffs are crossed by thrust planes that rise westward past a succession of coves and small headlands out to Whiten Head. Undulating cliffs run along the eastern shore of Loch Eriboll, parallel to the overthrust zone and fringed by quartzite. The loch is up to 110 m deep and flanked by shingle beaches, backed by steep slopes. West of Loch Eriboll is Lewisian gneiss cliffs that extend behind sandy bays. There is a deep cavern at the mouth of the Allt Smoo, a stream that descends into Smoo Cave, which is 60 m long and 40 m wide. The narrow promontory of Faraid Head has gentler western slopes descending to a wide sandy beach backed

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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7. 24. 6

North and North West Scotland (Duncansby Head to Cape Wrath and Fort William)

⊡⊡ Fig. 7.24.6.1 Geo in Old Red Sandstone near Dunscansby Head. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.6.2 Granitic Strathy Point, west of Pentland Firth. (Courtesy Department of Geography, University of Aberdeen.)

by dunes at Balnakiel. The peninsula of An Fharaid is a block of gneiss sloping south–west and bordered on its western side by the Kyle of Durness, cut back into limestones and quartzites (> Fig. 7.24.6.3). West of the Kyle of Durness, the north-facing coast has high cliffs cut at first in Lewisian gneiss and then in Torridonian Sandstone. Storm waves from the Arctic frequently break along this shore. The high cliffs towards Cape Wrath pass on to Torridonian Sandstone, and at Cape Wrath (> Fig. 7.24.6.4) the red sandstone cliffs are 110 m high, rising to the plateau of The Parbh.

3.  The West Coast Much of the spectacular coastline south from Cape Wrath lies west of the Moine Thrust. Lewisian gneiss forms the basement geology and has been glaciated to produce a very irregular and bare landscape. The gneiss is locally overlain by sub-horizontal beds of Torridonian sandstone that often terminate at the coast as steep and high cliffs. The only major beach in this area is at Sandwood Bay (> Fig. 7.24.6.5), which is over 1 km in length and backed by vegetated and semi-vegetated dunes. South of Sandwood Bay, the coastline is composed of Torridonian rocks with a conspicuous stack at Am Buachaille. The coast south of Loch Laxford to Enard Bay is mostly on gneiss, traversed in many areas by dykes of epidiorite. Locally, sandstone is exposed, as on Handa Island, where the coastal cliffs are up to 100 m high. At Scourie Bay, a coastal ridge of gravel fronts a small lagoon bordered by low ice-moulded gneiss.

North and North West Scotland (Duncansby Head to Cape Wrath and Fort William)

7. 24. 6

⊡⊡ Fig. 7.24.6.3 The Kyle of Durness is a sea loch that is almost entirely filled by silt, sand and gravel. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.6.4 Cliffs cut in Torridonian Sandstone at Cape Wrath, the NW corner of mainland Britain. (Courtesy Department of Geography, University of Aberdeen.)

South of Enard Bay, the coast is again dominated by sandstone, which dips at low angles to the west so that much of the coastline has inclined sandstone ridges sloping gently into the sea. In the area of Loch Broom and Little Loch Broom, numerous coastal areas display evidence of a Late Glacial 20–25 m sea level, such as the raised delta on which stands

the town of Ullapool. Locally, there are outcrops of Mesozoic strata, as in the low ground between Laide and Aultbea. To the south, Loch Ewe borders the NW extension of a large fault that trends along the northern shoreline of Loch Maree. The coastline is again dominated by gently inclined sandstone strata together with local stacks and geos. Across much of this area, bedrock is mantled by

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North and North West Scotland (Duncansby Head to Cape Wrath and Fort William)

⊡⊡ Fig. 7.24.6.5 Degraded dunes of Holocene age, Sandwood Bay, Sutherland. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.6.6 View of the Torridon hills showing inclined dip of rock strata. (Courtesy Jim Livingston.)

glacial drift associated with a Late Glacial ice advance, known as the Wester Ross Readvance. Sediment derived from glacial drift has supplied beaches as at Port Henderson, Opinan and Redpoint. To the south of Redpoint is Loch Torridon, type ­location for the Torridonian sandstone, which dominates the coast south to Applecross. Emerged beaches and platforms 20–28 m above sea level owe their origins to the

glacio-isostatic rebound following the melting of the last ice sheet (> Fig. 7.24.6.6). The coastline swings to the NE into Loch Carron at the southern limit of the Moine thrust. Loch Carron is strongly influenced by the distribution of gneiss and Torridonian sandstone. Near Strome Ferry, a ridge of gneiss almost divides Loch Carron into separate inner and outer lochs.

North and North West Scotland (Duncansby Head to Cape Wrath and Fort William)

The gneiss rocks outcrop south as far as Loch Hourn, with restricted areas of sandstone, schist and slate. There are numerous dolerite dykes on an irregular crenulate coastline cut through in places by glacial fiords. The largest of these are Loch Nevis and Loch Hourn, two of the deepest sea lochs in Scotland. South of Loch Nevis the coastline is increasingly dominated by schists, slates and phyllites. South of Mallaig, Loch Morar is separated from the sea by peat-covered ­glaciofluvial sands and gravels, raised beach and dune sediment deposited during the Holocene high stand of sea level. There are many beaches of calcareous white sand (Crofts and Mather 1972). South of Mallaig are some of the highest raised beaches in Scotland. Near Arisaig, raised marine terraces are found up to 42 m above Ordnance Datum. There are also slightly emerged shore platforms. The coastline becomes very irregular, with numerous sea lochs, at the head of which are gravel beaches backed by emerged Holocene beach gravels mantling glacial sands and gravels. Sand beaches are rare, while along parts of the coast are isolated raised sea caves. Loch Shiel is separated from the sea by an extensive area of glaciofluvial outwash gravels, from which are derived the beaches that flank Kentra Bay.

7. 24. 6

South of Kentra Bay is the Ardnamurchan Peninsula, the western part of which is a large Tertiary igneous complex. The coastal topography is complex with parts of southern and western Ardnamurchan dominated by ring dykes, cone sheets and volcanic lavas. The region of Sunart lacks glacial drift, and the coastline has few modern or raised beaches. Along the Sound of Mull, the schists give way to sheets of Tertiary basalt lava on a cliffed coastline. Along the western coastline of Loch Linnhe as far as Fort William, the coast drops steeply to the sea. The slopes are on granite and schist. The cliffs are interrupted at Cor­ ran by thick outwash gravels that extend SE across Loch Linnhe. They are limited seaward by an emerged cliff and terrace produced as a result of glacio-isostatic rebound.

References Crofts R, Mather AS (1972) The Beaches of Wester Ross. Department of Geography, University of Aberdeen, Aberdeen, Scotland Ritchie W, Mather AS (1969) The Beaches of Sutherland. Department of Geography, University of Aberdeen, Aberdeen, Scotland Ritchie W, Mather AS (1970) The Beaches of Caithness. Department of Geography, University of Aberdeen, Aberdeen, Scotland

503

7.24.7 The Inner Hebrides

Alastair Dawson

1.  Isle of Skye The Isle of Skye (>Fig. 7.24.7.1) in the Inner Hebrides (which are taken to include Skye and the islands to the south: Canna, Rhum, Eigg, Muck, Coll, Tiree, Mull, Colonsay, Oronsay, Jura and Islay) has an area  of over 1,600 sq. km and consists of a southern ridge, the Sleat Peninsula, of mainly Precambrian rocks, a central mountainous area of Tertiary intrusive rocks and a northern area of volcanic rocks. This account of the coast  of Skye begins at Kyleakin, by the Kyle of Lochalsh bridge and follows an anticlockwise sequence around the island. Between Kyleakin and Ob Lusa the coast is low-lying, and the gravelly shore exposed at low tide has outcrops of red sandstone. There are segments of a raised beach at 10–15 m above sea level, formed during a phase of higher sea level in the Late Pleistocene. At Ob Lusa, the Torridonian rocks give place to sandstones, limestones and shales. These form broad shore platforms, generally along bedding planes. The Ardnish Peninsula consists of limestones, and its northern coast has structural shore platforms with an irregular outline resulting from dissection along faults. There is a white beach of calcareous algal sand. To the north is Pabay Island, rising to 28 m, where limestones are overlain by shales. Its eastern, northern and western coasts are bordered by subhorizontal shore platforms, and there is a shelly beach on its northern coast. Shore platforms extend along Broadford Bay and round Rubh’ an Eireannaich, fronting a gently sloping coast. Scalpay Island is mountainous and consists largely of sandstone except for small patch of granite on its western shore; nearby is an emerged grassy cuspate foreland. The coast curves round into Loch Ainort, which is ­bordered by steep slopes in granite, and has a wide gravelly intertidal zone at its head, with patches of salt marsh. At the NE end is a small sandy cuspate foreland at Maol Ban. The coast then turns north, with granite being replaced by sandstone. The Meall a’ Mhaoil peninsula rises to 284 m, and on the southern shore of Loch Sligachan sandstone rocks form a slope behind alluvial Rubha Garbh. On the northern shore of Loch Sligachan the steep slope of An Leitir rises on Tertiary volcanic rocks to Ben Lee (445 m).

⊡⊡ Fig. 7.24.7.1 The Isle of Skye. T Trotternish Peninsula; W Waternish Peninsula; D Duirinish Peninsula; M Minginish Peninsula; S Sleat Peninsula; CH Cuillin Hills; RH Red Hills. 1 – Kyleakin, 2 – Pabay Island, 3 – Broadford, 4 – Loch Ainort, 5 – Loch Sligachan, 6 – Portree, 7 – The Storr, 8 – Rigg, 9 – Kilt Rock, 10 – Staffin Island, 11 – Kilmaluag Bay, 12 – Rubha Hunish, 13 – Lob Score Bay, 14 – Uig, 15 – Loch Losait, 16 – Waternish Point, 17 – Dunvegan, 18 – Neist Point, 19 – Bracadale, 20 – Drynoch, 21 – Talisker Bay, 22 – Loch Scavaig, 23 – Loch Slapin, 24 – Loch Eishort, 25 – Ord, 26 – Point of Sleat, 27 – Armadale, 28 – Ornsay Island. (Courtesy Geostudies.)

The slope is incised by numerous parallel stream gullies, between which are spurs mantled by glacial drift. The elongated island of Raasay has an area of Juras­ sic  rocks around granite hills in the south, Sandstone

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Inner Hebrides

f­ ormations occur in the centre with gneiss to the north. A narrow strait, Caol Rona, separate Raasay from the smaller gneiss island of Rona. Sections of the east coast of Raasay are composed of landslips beneath high basalt cliffs. North from Camas a Mhor Bay Jurassic rocks outcrop along the rocky shore, below younger igneous rocks. The coast steepens northward round Ben Tianavaig (413 m), and a vertical cliff (dolerite sill) 150 m high ascends northward from the coast behind steepland slipped terrain. As the cliff ascends, lava cliffs rise from a lower level to form a stepped topography with intervening grassy apron slopes along the east facing coast north as far as Udairn. At Udairn the coast turns west towards Portree. There has been extensive slumping on the slopes below the lava cliffs. East of Portree low terraces extend along the coast with outcrops of Jurassic rock formations (>Fig. 7.24.7.2). North of Portree a steep coast with lava cliffs extends to Holm, where the volcanic cliffs are fronted by a grassy talus apron. Holm Island (rising to 29 m) stands offshore, backed by a rocky reef with a gravelly intertidal isthmus in its lee. The abrupt coastal slopes continue north, before declining behind Bearreraig Bay. Slopes rise inland to The Storr, an upper lava cliff 720 m high near the Old Man of  Storr, a pinnacle that has slid away from the major basalt escarpment that runs from south to north, dominat­ ­ing the Trotternish Peninsula. The slopes are presently ­relatively stable, but there was extensive slumping under

periglacial conditions in Late Pleistocene to early Holocene times. The rock sequence is seen in the cliff at Kilt Rock (>Fig. 7.24.7.3). The upper sill recedes behind a sloping terrace below Garafad, and further north at Rubha Garbhaig a bold cliff of columnar basalt is fronted by talusmantled slopes. Cemented pebbles in a cave indicate two earlier phases of higher sea level (Richards 1969). At An Corran the coast curves westward into Staffin Bay. Staffin Island is a flat-topped sill, rising to 21 m, with a shore platform on its SE coast. Staffin Bay has a wide sand and shingle beach facing north, backed by a bluff incised by small streams. At low tide the Stenscholl River flows out over wrack-strewn boulders, as does River Brogaig to the west as coast curves round to Digg. Between Digg and Flodigarry the low cliffs and shore outcrops are backed by the irregular slopes of Quiraing, rising to the high volcanic escarpment that runs from The Storr. Terraces rise to The Table and The Needle, backed by a lava cliff that curves westward behind Loch Droighinn. At Flodigarry there are coastal landslides presently being undermined by wave erosion. Northwest from Creag na h-Eiginn the Jurassic rocks disappear, and are replaced by volcanic sill cliffs and shore platforms developed across beds of lava. Till-draped grassy bluffs and low basalt cliffs continue northward past Stac Lachlainn to Rubha Bheanachain, and on to Rubha nahAisaig, the NE point of Trotternish.

⊡⊡ Fig. 7.24.7.2 Scarp of Tertiary basalt rising above Middle Jurassic sandstone near Portree. (Courtesy Geostudies.)

The Inner Hebrides

7.24.7

⊡⊡ Fig. 7.24.7.3 Kilt Rock on the E coast of Trotternish, where an upper volcanic sill with columnar basalt 30 m high overlies horizontally stratified Middle Jurassic Valtos Sandstone, with a lower dolerite sill near sea level. (Courtesy Geostudies.)

A north-facing section of cliffs and coves extends westward past The Aird (111 m) to Rubha Hunish, a peninsula forming the NW point of Trotternish. The coastal slope declines to a wide shore platform in sandstone in Tulm Bay. The cliffs expose an olivine-rich sill with several alternating dark and light layers, while the beach to the south has a green colouration because of its locally-derived olivine content. To the south–west is the wide Lub Score Bay, with a sand and gravel beach backed by cliffs capped by a basalt sill. Further south cliffs of columnar basalt descend to the shore in Camas Mor Bay. Locally, sandstone rocks exhibit signs of contact metamorphism. The coast curves out to the hill known as the Stack of Skudiburgh. Cliffs rise to the Ru Idrigill headland, and then the coast turns eastward to Uig. High bluffs of volcanic rock descend to Jurassic outcrops on the shore. Uig Bay has a gravelly and bouldery shore exposed at low tide. The steep slope at South Cuil becomes cliffed towards Rubha Riadhain, where it is interrupted by Camas Beag and a small U-shaped bay on the headland of Ru Chorachan at the southern end of Uig Bay. The Trotternish Peninsula has a general westward dip corresponding with The Storr escarpment to the east. The River Haultin flows into Loch Eyre at the southern end of Loch Snizort Beag through a salt marsh and a wide area of seaweed-strewn gravel and boulders exposed at low tide. The coastal slopes have lava ledges and some low cliffed sectors and rocky shores, continuing round the Edinbane Peninsula and out to Lyndale Point.

To the west is Loch Greshornish, also bordered by lava slopes and minor cliffs that increase in size as one approaches Loch Snizort. The loch branches southwards to gravelly stream mouth inlets at Blackhill and Red Burn. Steep, low cliffs border the west coast of Greshornish Peninsula south to Loch Diubaig. The Waternish Peninsula runs out on the western side of Loch Snizort, with steep slopes and cliffs in Lower Tertiary volcanic rocks. There are natural arches in the cliffs, and rocky, boulder-strewn shores exposed at low tide. To the NW are the cliffs of Creag an Fhithich, declining westward to the low cliffed promontory of Waternish Point. The west coast, south from Waternish Point, has cliffs cut in volcanic rocks that are exposed to waves from the stormy Little Minch. They attain 120 m at Ard Beag, the northern point of the Ard Mor peninsula, which is attached to the mainland by a low isthmus at Trumpan. Ardmore Bay has a broad shingle beach behind sand. Jurassic rocks outcrop below the volcanic formations in cliffs at the head of the bay. The west coast of the Waternish Peninsula becomes lower and more subdued past Hallin, and near Lusta there are small outcrops of sandstone and shale. Glacial drift mantles the lower slopes, and has been cut back in the cliffs near Stein. On the western side of Loch Bay a relict landslide rises to the lava cliffs on Beinn Bhreac, as far as the Rubha Maol headland. Sedimentary rocks outcrop below the volcanic formations on the north-facing coast with a prominent

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The Inner Hebrides

dyke on Groban na Sgeire, where ridges of lava run out into the sea. The coast turns southward, behind Lampay Island, where there is a white beach of calcareous algal sand, gravel and shells. It is backed by low sandy machair with grassland and herbaceous vegetation, a feature common in Scotland but rare on Skye. To the south, along the eastern shore of Loch Dunvegan, a series of gravelly coves and minor headlands extend past Claigan to the crag on which Dunvegan Castle stands. Along the west coast of Loch Dunvegan slopes in volcanic formations descend to rocky ledges and boulders, with a beach of sand and gravel at Colbost. There is a natural arch at Am Famhar, and the coast steepens and becomes cliffy towards Dunvegan Head. South from Dunvegan Head, the coast facing the Little Minch is high and steep with lava cliffs between grassy slopes. Streams pour from the cliffs as waterfalls. The cliffed coast curves into the funnel-shaped bay of Loch Pooltiel, declining towards the head at Glendale, where there is a curved wave-built gravelly bank. On the south coast of Loch Pooltiel there is a small bay at Meanish (>Fig. 7.24.7.4). The coast steepens out to the cliffy An Ceannaich promontory, then turns south along the Milovaig Peninsula. High basalt cliffs run behind Oisgill Bay, and to the south a lava cliff curves inland behind the Neist Peninsula. The SW coast of Skye is exposed to strong wave action. East of the Neist Peninsula, Moonen Bay is backed by lava

sills and the bold scarp of Waterstein Head, with grassy landslides resting on volcanic formations. High almost vertical cliffs of columnar basalt occur here above bouldery screes along past Ramasaig Cliff and down to Ramasaig Bay. At Ramasaig Bay valleys converge on the coast to the south a high and steep rugged basalt coast continues to Idrigill Point. The coast then curves round along the western side of Loch Bracadale to Orbost. This is the rugged Durinish Peninsula, with glens cut by Pleistocene glacial meltwater and a steep coast incised by many streams. East of Orbost the shore of Loch Bharcasaig curves round to a cliffed headland with shore platforms on sandstones and limestones on the Greepe Peninsula. Emerged shore platforms border the Ardmore peninsula, which has small beaches of white coralline sand between bouldery shores exposed at low tide. The northern part of Loch Bracadale is a branching embayment fringed by gentle slopes. At Loch Harport an elongate inlet runs SE, bordered by broad gentle slopes on volcanic formations with only minor cliffs and a salt marsh at its head. To the south the coast rises westward to Rubha nan Clach point, steepening to vertical cliffs on the more exposed western coast. At Rubha Cruinn the cliffed coast turns eastward into Talisker Bay, with basalt cliffs diminishing to grassy valley-side bluffs. On the southern side an upland spur ends in the cliffed Leathad Beithe, with active screes, and the high cliffed coast to Talisker Point, below which is a large stack.

⊡⊡ Fig. 7.24.7.4 High cliffs cut in Tertiary basalt on the coast N of Meanish. (Courtesy Geostudies.)

The Inner Hebrides

South from of Talisker Point high cliffs with caves extend past Biod Ruadh to Sgeir Bheag, then decline as they curve in along the northern side of Loch Eynort. Along the southern side of Loch Eynort the coastal slopes are gentler, but they rise and steepen to Sgurr an Duine at the outer coast, and southward past Stac an Tuill, which has a natural arch. There are lava ledges, chutes and gullies on the steep coast that curves eastward into Loch Brittle, At the head this sea loch is Mussel Scalp, a gently shelving sand and gravel beach backed by low grassy dunes. On the southern coast of the loch there are cobble beaches at stream mouths below incised glens, and an emerged shore platform segment at Rubha na Creige Moire. At the Rubh’ An Dunain promontory the coast turns eastward along Soay Sound. Jurassic rocks appear at the base of the basalt cliffs, with some Cretaceous greensand and chert to the east. Shore platforms are poorly ­developed, and there is much kelp. The island of Soay is dominated by sandstone, forming steep cliffs on the western shore and bluffs descending to low cliffs on the east and south–east. The island is almost split by Soay Harbour on the north coast and wider Camas nan Gall, backed by a beach, to the south, both cut out along a major fault. Coastal slopes on Torridonian rock with Upper Jurassic formations at the base steepen eastward past a sandy beach east of Ulfhart Point. The sandstone outcrop ends south of the Eilean Reamhar tombolo, giving place to basalt on the west coast of Loch Scavaig, then very steep slopes of Cuillin

⊡⊡ Fig. 7.24.7.5 Sloping ramps of Jurassic sandstone on the coast south of Elgol. (Courtesy Geostudies.)

7.24.7

gabbro. The Camasunary Fault runs SSW across the coastline at the eastern end of the bay, bringing up a sequence of Jurassic formations along the west coast of the Strathaird Peninsula. The west coast of Strathaird Peninsula has lava cliffs over Upper Jurassic outcrops, and the shore to the south shows a succession of Jurassic formations dipping gently westward. Glen Scaladale is incised into a lava plateau and opens SW to Cladach a’Ghlinne, where calcareous sandstones and shales outcrop, At Elgol there is a gravel beach and cliffs and shore platforms with abrasion ramps. South of Port na Cullaidh sloping platforms on seaward-dipping sandstone are widespread (>Fig. 7.24.7.5). On the south coast of Strathaird Peninsula the Bearreraig Sandstone is underlain by sandstones and shales which form wide structural shore platforms backed by low cliffs. A broad strait leads into Loch Slapin and Loch Eishort. East of Loch Slapin the coastal slopes cross a series of geological Formations, including quartzites, sandstones and limestones. There are numerous dolerite dykes running SE–NW, some unweathered and protruding, others weathered and forming gullies. At Ord, quartzite rocks occur adjacent to a sandy beach. The coast is mainly formed in sandstone and schist. On the east coast of the Sleat Peninsula the Moine Thrust Zone runs across the peninsula of Ard Thurinish by the Aird of Sleat. Here, the schists pass beneath gneiss. In Armadale Bay a sandy beach runs along the northern shore, and there are wide shore platforms north to Knock Bay.

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The coast steepens on the sandstones north of Loch na Dal, with many small parallel streams flowing down to the Sound of Sleat. The Sound narrows into Glenelg Bay, and strong tidal currents flow through the steep-sided Kyle Rhea strait, which widens northward into Loch Alsh. As it does so, the Skye coast turns NW, then west, the steep sandstone slopes on running behind Loch na Beiste and westward to Kyleakin.

2.  Canna The isle of Canna is composed of basalt, with lava cliffs on the northern and western shores and dykes of basalt and some tuffs at the eastern end. Columnar basalt is seen spectacularly in the cliffs. Deep incised valleys descend to the north and south coasts. Tarbert Bay is the largest of several coves on the steep south coast, but there are few beaches. Sanday, a smaller island to the SE, also has cliffs of basalt. At low tide the two islands are linked by a broad sandy flat, with Canna Harbour on the eastern side and a bridge or causeway across the western end.

3.  Rhum The island of Rhum has much of its cliffed northern and eastern shore, including Loch Scresort, in sandstone. In Kilmory Bay the beaches are of derived brown sand, with layers and patches of white shelly sand and usually much seaweed. On the north shore at Samhnan Beach, a calcareous white sand is backed by machair and low calcareous dunes. There are many dykes in the coastal cliffs that continue round much of the island. Some gneiss crops out in the south, and there are sectors of gabbro and granite on the south–west and conglomerate and sandstone on the north–west coast. The western shore at Harris is bouldery below lava cliffs. A boulder beach impounds a small loch at Papadil in a valley between cliffs of gabbro.

4.  Eigg The island of Eigg is mainly Tertiary basalt but with intrusions of felsite in the high cliffs, bordered by apron slopes that descend to coastal cliffs. Similar cliffs and slopes surround An Sgurr (393 m), a prominent hill in the south of the island while the south west coast has steep slopes descending to cliffs. There are sandy beaches in coves around Laig.

North of the Bay of Laig are the “Singing Sands” at Camas Sgiotaig – well-sorted and well-rounded quartz sands that squeak when trampled and “moan” in the wind.

5.  Muck To the south of Eigg the island of Muck is entirely composed of basalt, with rocky shores and beaches of shell sand. Port Mor is in a south-eastern inlet, and Camas Mor is a southern sandy bay. The north coast has long narrow peninsulas and bays with a NNW trend.

6.  Coll The Island of Coll is low-lying, mainly craggy island developed in gneiss with strips of schist and foliated granite in the western part. There is a sector of cliffs on the east coast south of Sorisdale coast, but otherwise the coastline is irregular, with many coves. On the west coast these contain beaches of shell sand and calcareous dunes up to 30 m high, as at Hogh Bay which has a sandy beach backed by grassy dunes with blowouts. Raised beaches occur up to 25 m above sea level.

7.  Tiree The island of Tiree has rather featureless, low coasts with extensive pale sandy beaches fronting machair, and rocky shores with ribs of rock running out into the sea. It has shelly sand beaches backed by machair on both the NW and SE shores and cliffed backshore dunes. The island consists of gneiss with many dykes and patches of limestone on the sandy shore of Gott Bay. Mean spring tide range in Gott Bay is 3.5 m. There is a central isthmus developed in schist running south through to Hynish Bay. Only at the south-western end, on Ceann a Mhara, is there a steep headland, one of the three hills.

8.  Mull Separated from the Morvern mainland by the Sound of Mull, the large island of Mull has basalt dominating the coastline with lava plateaux and terraces separated by lava cliffs. The basalt is crossed by a swarm of dolerite dykes trending SE–NW, often forming headlands. In the SW there are outcrops of granite and sandstone with softer sedimentary rocks in the south and east.

The Inner Hebrides

Tobermory, at the northern end of the Sound of Mull, has a mean spring tide range of 4.0 m. To the north is Ardmore Point, with cliffs of basalt and a coast to the WSW incised by long narrow sea lochs. At Caliach Point the coast turns southward and Calgary has a sandy beach fronting a grassy machair at the head of a narrow bay. The basalt cliffs continue round Rubh ‘a Chaoil and border the Treshnish Isles offshore. The coast turns eastward along the northern side of Loch Tuach. The island of Ulva is separated only by a narrow strait, and consists of  lava flow terraces and gullies, occupied by extensive bracken. Only a narrow cleft separates similar Gometra island on its western side. Ulva also has one of the highest relict sea caves in Scotland at ca. 42 m above sea level. Loch na Keal is a wide sea loch bordered by steep slopes rising to mountains on the southern side. It contains basaltic islands, Eorsa and Inch Kenneth. Offshore is the island of Staffa, rising to 42 m and fringed by lava cliffs. It is famous for columnar basalts, including the structures at Fingal’s Cave (66 feet high, 228 feet deep) and the adjacent causeway at the southern end, Boat Cave and Clamshell Cave. South of Loch Na Keal, quartzite rocks run along the shore of the mountainous Ardmeanach peninsula, fringed by high cliffs. The coast curves round into Loch Scridain, another deep sea loch bordered by steep slopes in basalt, with the Coladoir River flowing in through a delta and intertidal gravels at its head. South of Loch Scridain the Ross of Mull runs out towards Iona, and has a low promontory of Eocene gravel at Carraig Chorrach. Mean spring tide range here is 3.7 m. Iona has sandstone in the eastern part and gneiss on the west coast with a small outcrop of schist in the south. A white sandy beach, Traigh Ban nam Monach, fringes grassy machair. Soa Island and other reefs to the south are composed of gneiss. The southern coast of Mull, east of the granite promontory, has cliffs cut in mica schist as well as small sandy bays. The high cliffs of Aoneadh Mor are probably of glacial origin, and in Carsaig Bay to the east there are outcrops of calcareous sandstone. Basalt forms the steep coasts bordering Loch Buie. Nearby, Lord Lovat’s Cave, 150 feet high, runs 300 feet into the cliffs. Beyond the mouth of the T-shaped Loch Spelve, Tertiary basalt returns to the coast past Grass Point and Loch Don, round to the dark headland known as Duart Point at the eastern end of Mull. Along the southern shores of the Sound of Mull there are outcrops of Conglomerate, sandstone and younger sedimentary rocks around Craignure Bay. The coast is

7.24.7

relatively low, with gravelly shores, particularly at valley mouths. Mean spring tide range here is 3.9 m. Basalt dominates the coast westward, as the coast steepens towards Tobermory.

9.  Colonsay and Oronsay Colonsay and Oronsay, out to the west of Jura, are islands of sandstone and mudstone, with numerous igneous dykes. The two islands are linked at low tide by a sand bar, The Strand, dries at low tide. There are many southern and western skerries. Kiloran Bay on the north–west coast of Colonsay has a sandy beach, washed by Atlantic surf and backed by dunes, cliffed along their seaward margins. There are caves in the sandstone cliffs at each end. Bordering peninsulas have platforms cut at higher sea levels. Scalasaig on the east coast has a mean spring tide range of 3.4 m. Oronsay has rocky shores.

10.  Jura South of the Gulf of Corryvreckan (which has strong tide races and whirlpools) quartzite, traversed by SE–NW dolerite dykes, dominates much of the island of Jura. The western coast is sloping, with several valley-mouth inlets, the largest of which is Loch Tarbert, which almost bisects the island. Glengarrisdale Bay has a mean spring tide range of 3.1 m. A 37 km sector south of Glengarrisdale Bay has emerged coastal landforms including shore platforms, shingle ridges and stranded cliffs. The emerged shingle beaches were related to Late Glacial and Holocene sea ­levels by McCann (1964). Emerged shore platforms include a High Rock Platform (rising to 32–34 m, a Main Rock Platform 3–5 m and an intertidal Low Rock Platform bearing remnants of glacial drift. There are also small pockets of beach, dune and machair, and a medial moraine at Sgriob na Caillich. The west coast is partly protected from Atlantic waves by the island of Colonsay, and south of Lake Tarbert it curves into the shelter of Islay. Cliffs cut in the quartzite are largely of glacial origin. Inland are three conical mountains, the Paps of Jura. At the south-eastern end of the Sound of Islay the low promontory of Rubha na Traille and the island of Brosdale are composed of schist. The gentle eastern coast of Jura, facing the Sound of Jura, has segments of schist between Lang and Rubha nan Crann, at Lussa Point, Ardlussa Bay and Barnhill.

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11.  Islay There are variations in mean spring tide range around Islay. Rubha a’Mhail at the northern end has .2 m, Ardnave Point 3.2 m, Orsay Island 2.2 m, Port Ellen 0.7 m, Port Askaig in the Sound of Islay 1.8 m. and Craighouse 1.0 m. The east coast of Islay, along the Sound of Islay, is also bordered by quartzite. Islay is a hilly peat-covered island with many streams. The quartzite rocks extend round the northernmost point, Rubha a Mhail, to north-facing cliffs interrupted by two areas of limestone. At Gortantaoid Point the low-lying coast, bordered by beaches and dunes that are generally stable or accreting curves into Loch Gruinart, where the dune fringe has been cut back at Killinanin. Loch Gruinart is fringed by salt marshes that overlie sand and gravel and stand behind wide sand flats exposed at low tide. The loch opens northward, and is 7 km long and 2 km wide, with a mean spring tide range of 3.1 m. Its shores are sheltered from strong wave action and the coast has been emerging. The salt marshes show terracing, and have narrow linear tidal creeks, with salt pans in the upper marsh. The head of Loch Gruinart is cut in sandstone which forms the western part of Islay, running out behind the dunes of Ardnave Point. The sandstone cliffs continue round to Machir Bay, which has a wide beach accreting behind intertidal sandflats, with calcareous sand still being

supplied to the shore. The beaches are backed by scarped marram grass dunes and machair. Indurated aeolian calcarenite has been exposed in blowouts through the grassy foredunes. The machair extends inland and is draped over a series of emerged marine terraces, rock outcrops, glacial drift and screes, ascending to about 60 m. A number of streams flow through the machair and dune fringe, which have been strongly influenced by hydrological processes. The south-western peninsula of Islay (Rhinns of Islay) has beaches washed by Atlantic swell and cliffs in gneiss while there are small bays in sandstone at Portnahaven, behind gneissic Orsay island, and a gneiss headland at Rhinns Point. The east coast then runs back along a curving fault (overthrust) into Loch Indaal, which is backed by sandstone mantled by glacial sediment. At low tide there is a wide sandy bay at the mouth of the River Sorn at Bridgend, and from Bowmore the sandstone extends round past Laggan Point. This has an 8 km long beach of shelly sand, the Big Strand. To the south of the Oa Peninsula there is a short fringe of basalt between Rubha Mor and Slugaide Glas, then cliffs in cut slates and schists extend south to the wild cliff of the Mull of Oa, where strips of limestone, conglomerate and quartzite are exposed along the steep southern coast (>Fig. 7.24.7.6). There are more slates and schists on Rubha nan Leacan, but Kilnaughton Bay at Port Ellen is cut back into a lowland on quartzite, and has a sandy beach. The quartzite runs along the irregular south-eastern shore to Ardbe and forms the island of Texa offshore. The

⊡⊡ Fig. 7.24.7.6 Cliffs bordering Mull of Oa, Islay, capped by Late Quaternary glaciomarine sediment.

The Inner Hebrides

irregular coast NE to Ardmore Point is fringed by many small islands and skerries, and has bays cut back in the softer rocks. North of Ardmore Point the east-facing coast has recurrent outcrops of schist and slates as far as Carraig Mhor, beyond Ardtalla, where quartzite returns to form the coast past McArthur’s Head into the Sound of Islay. The north coast of Islay has three distinct emerged shore platforms and associated raised shingle beaches. The High Rock Platform (about 33 m) is backed by an emerged cliff and mantled with glacial drift and marine deposits, and is thought to have been produced by cold climate

7.24.7

shore processes. Two younger rock platforms occur are close to present sea level while, nearby, Late Glacial emerged shingle beaches indicate formerly higher sea levels.

References McCann SB (1964) The raised beaches of north-east Islay and western Jura. Trans Inst Br Geogr 35:1–10 Richards (1969) Some aspects of the evolution of north-east Skye. Scot Geogr Mag 85:122

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7.24.8 The Outer Hebrides Alastair Dawson 

The Outer Hebrides (Westerln Isles) include the linked islands of Lewis and Harris, North Uist, Benbecula and South Uist, Eriksay, Barra and smaller islands south to Berneray (Angus 1997). The islands are dominated by PreCambrian rocks, (Lewisian gneiss), which is often irregularly layered. There are associated granites in the hilly uplands of ­western Lewis and Harris, some schist and limestone in ­southern Harris and a low-lying area of redbrown Permo-Triassic sandstones and conglomerates. The islands were strongly glaciated in Pleistocene times, and glacial drift blankets much of the northwest of Lewis, linking some former islands. Bold cliffs occur locally on exposed N and NW coasts, dissected along joint and fault planes, down which there have been local landslides, but undercutting and the retreat of headlands has been very slow on these hard rock formations, and even on the exposed north coast there are many vegetated slopes. The west coasts of the islands are exposed to Atlantic swell and storm waves, whereas the east coasts, facing the Minch, are exempt from swell but have occasional local storms. The prevailing winds are south-westerly: Stornoway has gales on an average of 48 days per year. Mean spring tide range is generally between 3.6 and 4.3 m (Stornoway 4.1 m, Bernera Harbour 3.8 m, Leverburgh 4.0 m, Barra 3.6 m, Loch Maddy 4.2 m, East Loch Tarbert 4.3 m). Beaches of sand and gravel derived from glacial drift occupy many bays, and there has been large scale accretion of calcareous sand, swept in by wave action from the sea floor, especially on the oceanic western coasts of the islands. Many beaches are backed by dunes and the lowlying calcareous sandy areas known as machair, some of which include subdued dunes and plains formed by deposition in shallow lochs (lakes) or lagoons (Ritchie and Mather 1977).

1.  Lewis and Harris This account begins at Stornoway Harbour, which is partly sheltered by the rocky Airinis peninsula. Stornoway ­waterfront is largely walled, with breakwaters enclosing

the ­harbour. There is a beach of sand and fine shingle in front of the sea wall and in the bay to the east. Tide gauge records show that mean sea level has been rising at Stornoway at about 2.5 ± 2.4 mm/year. The coast east of Stornoway consists of a series of small bays and rocky promontories out to Rubha Thuilm (Holm Point), the cliffs becoming bolder as exposure to The Minch increases. They are cut in reddish-brown sandstones and conglomerates which outcrop in cliffs east from here, extending round the Eye Peninsula and north along the west coast of Broad Bay. The cliffs are up to 10 m high, often slumping, with shore platforms that tend to follow the NW-NNW dip. Holm Bay, sheltered by the indented Holm Island, has rocky shores and reefs with brown seaweed. There are rounded grassy bluffs, bold, bull-nosed promontories and overhanging cliffs. Chubag Bay is wider, facing south, with cliffs cut in conglomerate (>Fig. 7.24.8.1), eroded to yield pink and grey pebbles that dominate beaches and tombolos. On the south coast of the Eye Peninsula grassy bluffs become higher and steeper, with a transition to bold rocky cliffs towards Chicken Head. In Phabail Bay there is an emerged shore platform cut in gneiss, much influenced by near-horizontal jointing. The rocky shore swings northward to Tiumpan Head, which has a prominent stack. On the northeast coast grassy slopes descend to a rocky shore with bands of hard dark gneiss and bouldery sections. Near Garrabost the Lewisian gneiss disappears beneath sedimentary rocks with grassy bluffs and fronted shore platforms cut in sandstones and conglomerates that dip westward. The harder conglomerates stand out as scarps running across the shore. North of Stornoway the coast of Broad Bay is fringed by a shingle beach, fronted by a broad sandy shore exposed at low tide. Laxdale Lagoon has a western salt marsh and enclosing spits. To the north is Coll Bay, where a sandy beach is backed by uncut dunes and fronted by pebble gravel. Cliffs and bluffs are in conglomerate, the upper part much weathered, the lower part forming a gravelstrewn shore platform. The promontory to Tolsta Head has vegetated bluffs passing east through slope-over-wall profiles to rocky

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Outer Hebrides

⊡⊡ Fig. 7.24.8.1 On the western side of Chubag Bay are cliffs and bluffs in Permo-Triassic conglomerate, with derived gravelly beaches and a shore platform. (Courtesy Geostudies.)

cliffs. Garry Beach (Port Geiraha) is a sandy valley-mouth cove backed by grassy dunes, the southern cliff fringed by several tall stacks cut in gneiss and shaped in relation to vertical and inclined joint planes. To the north of the coast is bold and rugged, with outcrops of pale gneiss. Beyond this the sedimentary rocks come to an end, and a rugged cliffy coast cut in gneiss runs past Cellar Head (>Fig. 7.24.8.2). To the north rocky headlands alternate with slumping bluffs cut in glacial drift. There are several caves and natural arches. Port Skigestra is a diminutive harbour on the NW side of a broad bay with boulders and a grey shingle beach at its head. The rocky shore ends in a sandy beach at Buile Muigh, near Port of Ness. To the north cliffs in black rock become higher with greater exposure to Atlantic storms. At the Butt of Lewis the headland consists of hard metasediment that have been dissected along joints and faults to form geos (chasms) and large stacks. The promontory consists of a 30–40 m platform. The more exposed cliffs have wider stripped zones exposing gneiss, with a set-back turf edge. To the south the natural arch through the headland at Roinn a’ Roidh is angled with the dipping structures. South of Europie, slopes with a thin mantle of wind-blown sand descend to rocky shores, then a wide beach of calcareous sand backed by dunes held by marram grass with some blowouts. At Suainebost is a long embayment with dissected grassy dunes behind a beach of sand and gravel. At Cros there is a cove at the mouth of the deep Dell valley, and a

stream flows through well-rounded cobbles and pebbles to a partly sandy beach, Traith Dhail. The cliff on the eastern side is cut across gneiss with landward inclined planes, and has an escarpment style. At Galson cliffs cut in glacial drift over dark gneiss include exposures of an emerged beach 3–8 m above high tide level. To the east the emerged beach is overlain by a shelly midden with charcoal and bone implements. To the south-west coastal slopes descend to low rocky cliffs, and there are bouldery shores and dunes in embayments. A beach of sand and gravel is capped by dunes along a barrier that encloses Loch Mor Bharabhais (Loch Barvas) (>Fig 7.24.8.3). This is backed by an almost flat machair plain. The beach runs on as a barrier across Loch Eirearaigh. Cliffs and coves continue SW to Port Arnol, a wider valley-mouth bay with a sandy beach. Cliffs and bluffs cut in glacial drift and underlying gneisses run past the Labost ridge round to Iuchair, the rocky headland beside the bay north of Shawbost. At Shawbost the steep coast descends to rocky cliffs and blocky shores in gneiss and a range of igneous and sedimentary rocks. Shingle beaches are derived from glacial drift gravels and the disintegration of shore outcrops. Loch Siabost is wide between parallel shores along which rocky headlands pass landward to cliffs and bluffs in bouldery glacial drift. On the western shore cliffs run out to the headland at Rubha Neidalt, bordered by a rocky cove between headlands stripped of glacial drift, the backing slopes on drift with disintegrating turf terracettes.

The Outer Hebrides

7. 24. 8

⊡⊡ Fig. 7.24.8.2 Steep rocky coast north of Port Geiraha, looking towards Cellar Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.24.8.3 Gravel barrier enclosing Loch Barvas on the west coast of the Isle of Lewis, with low cliffs cut in glacial drift. (Courtesy Geostudies.)

To the south-west is Loch na Muilne, a large cauldron linked by a tunnel to deep geos, with waves swashing in to a shingle beach piled with driftwood. On Rubha na Beirghe is a large natural arch through a stack. The cliffs and coves continue SW to Dalbeg (Dhailbeag) Bay, with a group of stacks and a beach of calcareous sand receiving Atlantic swell. In general, valley-mouth bays on the NW coast of Lewis have received inwashed calcareous sand whereas rocky and bouldery coves have not, proabably

because they reflect strong wave action and generate seaward sweeping of such sediment. To the west of Balbeg Bay is a broad, high rocky ridge, with ice-moulded and scoured knolls and boggy hollows, some with lochans. The ridge is truncated at the coast by high rugged cliffs, with a natural arch out on the western headland. Cliffs run out SW to Aird Laimisiadair and the entrance to Loch Rog an Ear, a deep inlet formed by partial marine submergence of a glaciated valley. To the south

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The Outer Hebrides

is Loch Shiadan, a gorge partly blocked by a boulder ridge. An irregular shore platform follows sloping structures in the gneiss. At Breascleit a cobble beach fronts a cliff cut in glacial drift. Spring tide range here is about 3 m. To the west Great Bernera has an indented coastline, with inlets and peninsulas that follow a NNE-SSW trend. At the northern end Bosta has a sandy beach in a valley-mouth bay bordered by slopes with a mantle of wind-blown sand. Erosion during a 1996 storm exposed an Iron Age village. West of Great Bernera is Loch Rog, a wide branching marine inlet which penetrates several kilometres into a landscape of ice-scraped, ice-moulded rocky hills. The Valtos Peninsula has rocky southern and eastern shores and calcareous sand on parts of the northern coast, notably in the long gently curving beach Traigh na Beirigh. Wave action is impeded by offshore islands, but swell moving in through a strait (Caolas na Sgeire Leithe) is refracted into a curved pattern that shapes the beach. On the western side of the Tobha Mor promontory Cliobh Beach is wide and sandy,and the cliffed coast runs on NW to Gallan Head. Cliffs south from Gallan Head, exposed to Atlantic storms, have wide stripped edges exposing bare rock. The broad bay of Camas Uig is backed by the wide sands of Uig Bay, Traith Uig. The sand has come partly from erosion of a nearby esker and is partly calcareous, swept in from the sea floor. Some of the inner parts of Uig Bay have patches of grassy salt marsh with winding creeks (>Fig. 7.24.8.4).

West of Camas Uig Bay the cliffs show little evidence of recession apart from local rock falls, and there are irregular rocky shores rather than shore platforms. To the south is Mangerstra, which has high cliffs behind a deep cove with boulders. At the cliff crest turf strips are breaking and collapsing along fractures parallel to the coastline, and the pasture above the cliff top is sandy machair. Out to sea is St. Kilda (Hansom 2003), the largest of a group of outlying Hebridean islets and stacks, and is what remains of a large Tertiary ring volcano, its steep cliffs rising directly from a 40 m deep submarine platform which itself is fronted by steep submarine cliffs that rise from an even deeper submarine platform at 120 m. Caves have been cut into both the submerged and the modern cliffs but modern equivalents of the submerged platforms are absent. Hirta is the largest island (630 ha), with Dun, Soay and Boreray comprising the 853 ha land area of this World Heritage Site. The islands were subject to mountain glaciation during the last glacial maximum but parts of Hirta were ice free and are mantled with periglacial material. Spectacular, often vertical and overhanging, cliffs surround all the islands of the group, Conachair (430 m) on Hirta being the highest sea cliff in Britain. Many of the cliffs pass upwards into steep vegetated coastal slopes (>Fig. 7.24.8.5). Stacks are common and close to Boreray, Stac an Armin and Stac Lee tower 191 and 166 m above sea level, the former being the highest stack in the British Isles. Rockall, 300 km to the west, is an isolated stack of bare granitic rock 21 m high, standing on a submerged

⊡⊡ Fig. 7.24.8.4 Dissected salt marsh behind Uig Sands, NW Isle of Lewis. (Courtesy Geostudies.)

The Outer Hebrides

⊡⊡ Fig. 7.24.8.5 Looking east along the south coast of Hirta, the main island of the St Kilda group. (Courtesy J. Hansom.)

⊡⊡ Fig. 7.24.8.6 Sanded inlet at Mangerstra, NW Isle of Lewis. (Courtesy Geostudies.)

7. 24. 8

basalt bank. In February 2000, the highest individual wave ever recorded (29.1 m) was experienced near Rockall. Further south is Mangerstra Inlet, a wide valley-mouth bay with upper sand flats bordered by grassy dunes formed by deflation down to the water table (>Fig. 7.24.8.6). The sand flats are occasionally invaded by the sea when high spring tides are augmented by onshore winds and incoming waves. The west-facing coast south of Mangerstra is backed by a wide gently sloping terrace, steepening to rocky mountains. The cliffs are irregular, some gnarled, others with vertical or steep joint planes, some dissected into stacks. Offshore is the island of Eilean Mhealasta, with a beach on its eastern coast, and the higher island of Scarp beyond. The mountainous west coast continues southward, broken by sea lochs. At Huishinish Bay a south-facing beach of shelly sand fronts stable machair. The coast then runs ESE into West Loch Tarbert, which narrows to the isthmus on which the little town of Tarbert stands. The southern peninsula of Harris is also mountainous, the NW coast partly sheltered from ocean waves by the high island of Taransay, which has a cuspate spit (Corran Ra) on its inner shore, facing the Sound of Taransay. A wide estuary at Luskentyre contains inwashed sand, which is exposed as a broad plain at low tide. The upper estuary has brown silty sand, passing downstream to yellow inwashed calcareous sand, and the southern shore has rocky coves with small sandy beaches. The broad Corran Seilebost sandspit is dune-capped, the dunes cliffed on both its inner and outer shores.

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The Outer Hebrides

Southwest from Luskentyre there are several sandy beaches on islands separated by steep rocky sectors that have been stripped of turf by storm waves and ramps following low-angle joints. Sandflats run back on the eastern side of the Magillivray Ridge to a shallow lagoon and a much-dissected sandy salt marsh. On the SW side of Scarasta Sands is a hilly isthmus south of the high Toe Head mountain, on the slopes of which a pale pegmatite dyke slopes gently eastward. There is a natural arch out on the headland and sandy coves on the southern side of the mountain, then a broad sandy beach (Traigh na Cleabhaig) to the SW, backed by erod­ ­ed  and dissected grassy dunes, with a gently declining machair to salt marsh. At Leverburgh the SW coast of Harris shore has rocky shores with a prominent black algal horizon. Killegray and Ensay are steep-sided high islands in the Sound of Harris, and to the west is the conical island of Pabbay. To the SE is Renish Point, the southernmost point of Harris. The southeast coast of Harris is rocky and indented, with many inlets between irregular low promontories and few beaches. Aird Mhighe is a tidal pond (>Fig. 7.24.8.7), with water streaming out as the tide falls and returning as it rises. The rugged indented coastline of the east coast of Lewis, bordering the Minch, is dissected by several large sea lochs, notably Loch Seaforth, a straight fiord with steep slopes cut across the Hebrides Thrust near its entrance. Steep, rounded slopes descend to rocky shores along The Minch northward to Stornoway.

2.  North Uist On the southern side of the Sound of Harris is the island of Berneray, its eastern half high and hilly, while to the west is a machair lowland fringed by dunes and sandy beaches. The vast mass of deposited sand on the west coast of the Uists comes partly from sandy moraines on the Uists, extending below sea level, but the calcareous fraction is commonly at least 40%. There are salt marshes at Port nan Long, and to the south of the hill at Sudanais a wide machair with dunes and a sandy beach, Traith Lingeigh, on the seaward fringe. There are areas of dissected salt marsh and eroded peat along the southern shores. A tombolo, Machair Leathann, links a hilly peninsula with extensive dunes. Traigh Iar and Traigh Bhalaigh are long curving sandy beaches shaped by ocean swell. The NW coast of North Uist is rocky with small bays and rocky promontories. A beach backed by dunes curves out to a cuspate point, Rubha Mor, in the lee of small rocky islands, and Lagan Arnal, to the south, marks the beginning of the long sandy west coast of North Uist. Kirkibost Island and Baleshare have the form of barrier islands fronting a wide lagoon at high tide, backed by an irregular inner shoreline. At low tide wide sandy and muddy flats are exposed, with tidal channels that converge to gaps between the islands. Kirkibost Island is a dunecapped sandy barrier island between tidal entrances to the broad lagoon, its inner part low-lying machair.

⊡⊡ Fig. 7.24.8.7 Tidal pond at Aird Shleibhe, Harris, showing the outflow through a rocky channel as the tide falls. (Courtesy Geostudies.)

The Outer Hebrides

7. 24. 8

⊡⊡ Fig. 7.24.8.8 Cliff cut into machair on the west coast of South Uist near Verran. (Courtesy Geostudies.)

The west coast dunes on Baleshare Island are backed by machair, then a low area of glacial drift and rocky outcrops. The island is attached to the mainland by a causeway with sand accreting on the northern side. The east coast of North Uist, facing The Minch, is rugged, rising to rocky ridges that are penetrated by several sea lochs, notably Loch Euphoirt, which almost divides the island. In contrast with the oceanic west coast there are very few beaches. To the north is Loch Maddy with many seaweed-fringed islets.

3.  Benbecula Causeways link North Uist to the island of Benbecula at Gramsdale. On the western side is An Tom, a dune fringed machair peninsula. This low-lying terrain has been used for Ballyvanich airport. Dunes back the broad beaches on the SW coast. The east coast of Benbecula is, like that of North Uist, rocky and indented, with many small islands and few beaches.

4.  South Uist Much of the western half of South Uist is low-lying machair bordered seaward by a dune fringe. The ocean coastline is receding, and some settlements have been lost. Breaching of the dune fringe has exposed parts of the machair to erosion (>Fig. 7.24.8.8). There are successive promontories and sandy bays with lobate protrusions in the lee of rocky reefs.

The coast swings eastward at Pollachar, and faces south across the Sound of Barra. The high island of Eriksay is linked by a causeway built in 1998, and has a west coast bay with a sandy beach known as Prince Charles Beach. The island is almost divided by a rock-edged inlet, Acairseid Mhor. The east coast of South Uist, facing the Little Minch, has generally steep slopes descending to rocky shores which run in around Loch Boisdale, with its many islands, Loch Eynort and Loch Skipport. Again there is a contrast with the oceanic western coast in that sandy beaches and dunes are absent.

5.  Barra Much of the island of Barra is bare and mountainous, consisting mainly of Lewisian gneiss scoured by glaciation. The coast is sloping and rocky, indented by many small bays, but in the north the sandy isthmus of Eoligar­ ­r y links a hilly northern area. On the western side of this isthmus is Traigh Eais, a sandy beach that is wide at low tide, with some pebbly areas, backed by high cliffed grassy dunes, then a broad low undulating machair plain. On the eastern side is Traigh Mhor, a broad sandflat exposed at low tide which is a notable cockle resource. Splays of shelly sand are moved across the sandflat by wave action at high tide, and are harvested for shell grit. Aircraft land on the Traigh Mhor sandflat at low tide (>Fig. 7.24.8.9). Shelly sand is still moving in from the north through the Sound of Orosay.

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The Outer Hebrides

⊡⊡ Fig. 7.24.8.9 An aircraft landing on the beach at Traigh Mhor, Barra. (Courtesy Geostudies.)

To the south of Barra is the high island of Vatersay, linked by a causeway. It has a central isthmus consisting of beach-fringed parabolic dunes and machair between Bagh Siar and Vasteray Bay. As on Barra, the existence of an east coast sandy beach is probably the result of overwashing from the western side before the dune-covered isthmus developed. Other glaciated mountainous islands southward are Sandray, Parbay, Mingulay and Berneray, each with steep coastal slopes descending to cliffs and rocky shores, but only very small beaches.

References Angus S (1997) The Outer Hebrides. The shaping of the islands. White Horse Press, Cambridge, UK Hansom JD (2003) St Kilda, Western Isles. In: May VJ, Hansom JD (eds) Coastal Geomorphology of Great Britain. Geological Conservation Review Series No. 28, Joint Nature Conservation Committee, pp Peterborough, Cambridgeshire, 60–68 Ritchie W, Mather A (1977) The Beaches of the Highlands and Islands of Scotland. Countryside Commission of Scotland, Perth, Australia

7.24.9 Fort William to the Clyde Alastair Dawson 

This part of the west coast of Scotland consists of elongated peninsulas and islands separated by deep, narrow lochs, all trending NE–SW (Steers 1973). Mean spring tide ranges are Port Appin 2.8 m, Dunstaffnage Bay 3.3 m, Oban 3.3 m, Seil Sound 2.4 m, Gigha Sound 1.1 m, Southend, Kintyre 1.9 m, Campbelltown 2.15 m, Inveraray 3.1 m, Rothesay Bay 3.1 m, Lochgoilhead 3.0 m, Arrochar 3.1 m, Gairlochhead 3.1 m, Helensburgh 3.0 m. From Fort William the straight steep coast of Loch Linnhe runs SW below mountains that rise to Ben Nevis (1,344 m). Slopes on drift-covered Moine schists descend steeply to the coast. The coastline is guided by the Great Glen fault that extends SW as far as the Isle of Islay. This coastal area was glaciated, so that the only emerged shorelines are those produced during the Holocene. They are typically below 15 m OD, and are extensive in the Fort William area. There are occasional gravel beaches, as at Coruanan, Bunree and Onich, and Loch Leven is a deep narrow loch that comes in from the east (> Fig. 7.24.9.1). At Ballachulish there are slate beaches derived from waste from the slate quarries dumped on the loch shore.

Between Loch Leven and Oban the coastline follows fault lines. There are sectors of glacial outwash gravel and associated Holocene raised beaches. The elevation of these varies regionally as a result of differential glacio-isostatic recovery. The peninsula south to Keil has strips of limestone with quartzite on the point. Shuna Island is of slates and schists in the north and limestone in the south. The southeastern shores of Loch Linnhe are fringed by slates and schists and there is an emerged bench in front of bluffs that were formerly cliffs. The Port Appin peninsula is also cut in quartzite, slates and schists, with a bay cut in limestone before the coast curves round into Loch Creran. The elongated island of Lismore, in the SW of Loch Linne is a gently undulating island of limestone traversed by many dykes. Loch Creran has boulders, cobbles and rocky shores. Sandstone extends on to the spur to the south. At the narrow entrance to Loch Etive are the Falls of Lora, where as the tide falls the loch water pours out over a rock ledge into Loch Linnhe. Loch Etive is a sea loch bordered by glaciated slopes of sandstone.

⊡⊡ Fig. 7.24.9.1 Ballachulish, with a beach of slate quarry waste. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Fort William to the Clyde

Sandstone also borders much of the Oban coast. The Main Rock Platform is well developed here, the lower part of the town of Oban built on it. It is backed by an emerged cliff line, with the Dogstone representing an uplifted sea stack. The platform altitude decreases W and SW from Oban at a regional gradient of 0.14 m/km until it passes below present sea level (> Fig. 7.24.9.2). The Sound of Kerrera is bordered by raised beaches and the island of Kerrera has sandstone cliffs on the western shore, and schist and slate to the east. Loch Feochan is cut into lavas. Clachan Bridge, built in 1792, crosses a very narrow N–S channel (Seil Sound) to the island of Seil, which has cliffs on its west coast. Seil Sound has outcrops of black slate, phyllite and mica schist. It widens southward past the islands of Luing, Torsa and Shuna. There are strong tidal currents between the islands. The Garvellachs are islands of tillite and the cliffed western and southern shores of Scarba are in quartzite, which forms craggy outcrops. To the south is the Gulf of Corryvreckan, which has strong tides and whirlpools, and beyond that the island of Jura. The mainland coast curves round to Loch Melfort between wide gentle Slopes. The loch has grassy marsh and shores of gravelly mud shores. Cliffs cut in slates and phyllites and shingle beaches dominate the shores of the Craignish Peninsula, and to the south of is Loch Craignish, also trending SW. Crinan Bay is set back, with low wooded cliffs, and the end of the canal from Lochgilphead over to Loch Fyne to

the east. Carsaig Bay is sandy and backed by a narrow isthmus across to Tayvallich. There has been dissection along the SW–NE strike, forming peninsulas, islands and inlets in that direction. On the eastern side Loch Sween runs NE into Loch a’ Bhealaich, which has elongated peninsulas and straits following the geological strike of the country. A ridge ends in the cliffy Point of Knap, where the coast turns NE beside Loch Caolisport. Slates and schists with intruded dykes fringe both shores of Loch Caolisport (> Fig. 7.24.9.3). There are sandy beaches with shells and gravel, outcrops of sandstone behind wind low tide flats. Dalradian slates and schists outcrop on Kilberry Head. The coast turns east into Loch Stornoway, where there is a gravelly raised beach, occasional spurs and some segments of rocky cliff. The northern shore of West Loch Tarbert is cut in quartzite. Limestone emerges towards the head of this loch, and mica schist runs along the southern shore, the start of the steeper coast of the Kintyre Peninsula. Gigha Island has cliffs cut in schist, extending through the western half, quartzite in the eastern, crossed by dykes, with schist at the northern end. Cara Island is also developed in quartzite rising to a southern high point at the Mull of Cara. Glenacardoch Point is a triangular cliffed promontory of Dalradian quartzose mica schist, and there are sectors of Old Red Sandstone to the south and along Bellochantuy Bay (> Fig. 7.24.9.4). An emerged shore platform is backed

⊡⊡ Fig. 7.24.9.2 Main Rock Platform and cliff at Kerrera, near Oban. The inner edge of the platform occurs here at about 10 m OD. A distinctive characteristic of this platform and cliff throughout the Inner Hebrides is that the features often occur in areas of sheltered fetch.

Fort William to the Clyde

7.24.9

⊡⊡ Fig. 7.24.9.3 Loch Caolisport, showing planed-off slate and schist on the shore. (Courtesy Geostudies.)

⊡⊡ Fig. 7.24.9.4 Bluffs mark an emerged coastline in Bellochantuy Bay. (Courtesy Geostudies.)

by bluffs that become bolder with increased exposure to Atlantic waves between Islay and Ulster. Macrihanish Bay has over 3 miles of pale yellow sand, and has been shaped by Atlantic swell arriving through the gap between Ireland and Islay. South of Macrihanish the coast is high and steep, cut in mica schists, extending down to the Mull of Kintyre. On the Mull of Kintyre a broad bench is backed by a scarp in gnarled grey rock, and fringed by a bouldery shore The steep coast continues east to Carnsey Bay and Dunaverty Bay, which has a sandy beach backed by dunes,

with ribs of rock running out across the shore. This part of the coast is sheltered from Atlantic waves by Ireland, only 24 km away. Sandstone forms the steep coast from Polliwilline Bay round to Johnstons Point and Achinhoan Head. The coast faces eastward across Kilbrannan Sound to the island of Arran. Along this coast are exposures of pink glacial drift above sandy beaches, some steep headlands, little glens, bays with sand and gravel beaches between rocky spurs of schist. Campbeltown Loch is cut in schist, with high, bare hills behind.

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Fort William to the Clyde

North of Campbelltown there are outcrops of limestone on either side of sandstone in Ardnacross Bay. North to Saddell Bay on the east coast of the Kintyre peninsula is cut in schist, with a dolerite dyke at Carradale Point.

1.  Arran Offshore from Claonaig Bay, across where the Sound of Bute meets Kilbrannan Sound, is the Island of Arran. It has a large granite intrusion which does not reach the coast; there are several geological formations in the coastal fringe. The northern part is mountainous, the southern a lower plateau. The northern point, Cock of Arran, is on sandstone, with a segment of limestone to the west. To the west Lochranza is a little sea loch at the mouth of a deep valley, with a wide pebbly beach. From Lochranza to Dougarie schist forms the steep NW coast, curving round the margins of the granitic intrusion. This gives place to sandstone, which is interrupted by a narrow strip of lava south of Dougarie. There are a few sandy beaches. The sandstone comes out behind Machrie Bay. Cliffs extend past Kings Cave south to Drumadoon Point. Lava outcrops form the bold SW coast on either side of sandstone and conglomerate at Blackwaterfoot, which has sandy beaches. There are several dykes on Brown Head. The coast curves eastward in sandstone and conglomerate with an interruption of lava at Bennan Head. Pladda Island off the south coast is of basalt. The east coast is sandstone with intrusions of basalt making headlands, as  at Dippen Head, Larrybeg Point, Whiting Bay and Kingscross Point. Holy Island, off Lamlach Bay, is of sandstone north and trachyte in the south, and rises to 314 m. Coastal slopes descend to rocky outcrops behind a rugged shore littered with boulders. On the east coast of Arran Lamlash Bay has a shingle beach. Basalt outcrops on Clauchlands Point. There has been sand accretion in Brodick Bay, backed by a steep rise to Goat Fell (874 m). To the north sandstone gives place past Corrie to a south-to-north sequence of narrow outcrops of coal, limestone and sandstone along the coast.

2.  Carradale Point to the Clyde On the east coast of the Kintyre Peninsula north from Ca­­ rradale Point Slopes descend to low cliffs. From ­Cla­onaig

Bay past Skipness Point the cliffs are cut in mica-schist, with many dykes of basalt intersecting the coastline, some forming small promontories. Skipness has a curving sandy beach. From Skipness Point the steep coast runs northward along the western side of Loch Fyne to East Loch Tarbert. A narrow isthmus separates Western Loch Tarbert, which runs SW into the Sound of Jura. North of Tarbert are outcrops of schist with strips of limestone. Barmore Island is attached to the mainland by a narrow isthmus and Glen Shira opens into a small tributary bay, Loch Shira. The western shore of Loch Fyne northward is cut in quartzite. The beaches are of grey glacifluvial shingle. To the north is Ardrishaig with a gravel beach, a little harbour, and a lock at the mouth of the 14 km Crinan Canal cut by John Rennie in 1794 across to the Sound of Jura. Loch Gilp is cut in limestone. At Lochgilphead is a bay that at low tide becomes a wide shoal of muddy sand. More quartzite comes in towards Loch Gair. On the NW shore Inveraray is in slates and phyllites. The head of Loch Fyne at Clachan (Glen Fyne) has grassy marshes. Between Loch Fyne and Loch Long is the Cowal peninsula. The eastern shores of Loch Fyne run across successive formations that trend NE–SW, Down to Ardlamont Point the indented cliffy coast continues along the eastern shore of Loch Fyne. There are sandy beaches in Asgog Bay while accretion by shoreward drifting has taken place in Kilbride Bay. Inchmarnock Island to the south is in schist, with cliffs on the western coast. The western shore of the Kyles of Bute is bordered by schist, forming the peninsula north from the cliffs of Ardlamont Point. The island of Bute has similar coastal features, with a sandy beach in Ettrick Bay. Its southern end is low-lying on sandstone, with a cliffy peninsula of basalt extending out to Garroch Head. At Rothesay an esplanade backs a little curving sandy beach. Loch Striven borders another mountainous peninsula that declines to a south-facing coast between Ardyne Point and Toward Point. North of Toward Point the steep coast of sandstone descends to the Firth of Clyde. Dunoon is a seaside resort, and to the north a steep coast continues past Holy Loch to Loch Long and its branch, Loch Goil. On the high, grey Ardgoil peninsula to the west, Beinn Reith and Saddle Peaks reach over 600 m. On the eastern shore of Loch Long, there are localised beaches of gravel. At Kilcreggan the coast curves round into the Gare Loch. Along the east coast of Gare Loch sandstone extends to Rhu and then past Helensburgh.

Fort William to the Clyde

At Cardross the sandstone is cut by a major fault. On the north shore of the Clyde this sandstone extends from Dumbarton (on a crag of basalt over what was the lowest fordable point on the Clyde) east to Clydebank, across the river from Glasgow.

7.24.9

Reference Steers JA (1973) The Coastline of Scotland. Cambridge University Press, Cambridge, London

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7.24.10 South-West Scotland

William Ritchie

1.  The Clyde to Loch Ryan Downstream along the River Clyde from Glasgow, except for Dunbarton rock, the estuarine coastline on both sides is low with discontinuous towns and industries. At Gourock the coast swings SW past Cloch Point. South from Inverkip the coast borders wooded hills. South of Wemyss Bay the cliffs rise to wooded slopes. A grav­ elly  beach extends to Largs, and red stones continue past Fairlie. Offshore is the Cumbraes which are small islands composed mainly volcanic rock outcrops. There are numerous intrusive dykes, as on Farland Head and sandstones dominate the coast SW from Ardnell Bay to Ardrossan. Saltcoats has a harbour with a breakwater, and fossilised trees are exposed at low tide in a sandy intertidal zone. Irvine Bay has a long curving sandy beach, backed by dunes (>Fig. 7.24.10.1) and extending past the mouth of the River Garnock. At Troon sand drifting southward has accumulated to widen the beach in North Sands, on the north side of this

peninsula. While South Beach is narrower and has been depleted. A sandy bay curves south to Prestwick, where there are outcrops of dolerite. Ayr is a former port at the mouth of the River Ayr with sandy beaches. To the south cliffs rise to the Heads of Ayr. The cliffs rise south of Dunure. Seg­ ganwell Gorge is cut out of softer sandstone. Lavas recur between Heads of Ayr and Culzean Bay (>Fig. 7.24.10.2). At Port Carrick some sand accretion has occurred and a sandy beach, backed by dunes, curves southward round Maidenhead Bay. At Turnberry there is a long beach of fine silver sand backed by grassy dunes. Sandstones dominate the bluffs that runs behind the coast south from Turnberry, passing to Ordovician rocks. Off Girvan is Ailsa Craig, 338 m high and 3 km in circumference, the core of an ancient volcano (>Fig. 7.24.10.3). Little Ailsa to the SW has numerous vertical basalt pillars. At Girvan a sandy beach has segments of irregular shore platform cut across dipping strata. Ardmillan Castle stands on a coastal terrace backed by steep bluffs.

⊡⊡ Fig. 7.24.10.1 Between Saltcoats and Ayr the coastline is a nearly continuous sweep of sandy beaches and dunes. (Courtesy Department of Geography, University of Aberdeen.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.24.10 © Springer Science+Business Media B.V. 2010 (Dordrecht)

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South-West Scotland

⊡⊡ Fig. 7.24.10.2 North of Turnberry Devonian lavas outcrop to form a prominent rugged topography. (Courtesy Department of Geography, University of Aberdeen.)

⊡⊡ Fig. 7.24.10.3 From Lendalfoot on a coastline of volcanic rocks there is a view out to Ailsa Craig, a volcanic plug in the centre of the wide Firth between Kintyre and the Ayrshire coast. (Courtesy Department of Geography, University of Aberdeen.)

South of Ardwell Bay is a steep coast on ultrabasic rocks behind tuffs. Lendalfoot has a small shingle beach with gnarled rocky outcrops, backed by some dunes and a coastal terrace on which stand emerged rocky stacks. Bennane Head is cut in grey desert sandstone. The bluff runs out to a high rocky cliff of gnarled grey rock, with caves in an emerged cliff. South from Bennane Head a sand and shingle beach extends to Ballantrae, where a

high grey shingle beach deflects the mouth of the River Stinchar southward. From Ballantrae to Currarie the steep coast is fringed by basalt, and to the south Glen App follows the line of the Southern Uplands Fault. Old House Point and Cairn Point are low cuspate protrusions into Loch Ryan. Along the eastern side of Loch Ryan there are gentler slopes in a variety of sedimentary and volcanic strata.

South-West Scotland

2.  Loch Ryan Cliffs on the SE shore of Loch Ryan become bluffs that run on inland as the coast curves to the north-facing shore of the isthmus south of Loch Ryan. At the SW end of is Stranraer where the north-facing sandy beach is very wide at low tide. A sandstone bluff of Permian Desert Sandstone runs north to Kirkcolm. There are wide intertidal mudflats with some shells, sand, and scattered boulders. The Wig is a shallow bay backed by a shingle beach and the coast curves out to The Scar, which is probably related to a glacial moraine. North of Kirkcolm the cliffs pass into wooded bluffs fronted by dissected shore ledges of red sandstone breccia, with segments of pebble beach. Lady Bay has a beach of sand and backing shingle, but there are extensive intertidal rock outcrops. At Milleur Point, beside the entrance to Loch Ryan coastal slopes descend to low cliffs and rocky ribs cut in metamorphic outcrops.

3.  The Rhinns of Galloway The long narrow Rhinns peninsula has shore outcrops of mainly metamorphic rock types some hard and protruding, others excavated as shore clefts. Cliffs are generally high along the steep west coast, descending to rocky shores where at low tides, kelp is visible, attached to rocks that are almost always submerged. The coastal slopes are in glacial

⊡⊡ Fig. 7.24.10.4 Dissected coastal terrace south of Dally Bay. (Courtesy Geostudies.)

7.24.10

drift over solid rocks, and the grassy slopes have chutes, cauldrons, dells and fluting in the drift mantle, possibly produced by melting snow in periglacial phases of the late Quaternary. Beaches are of local derivation, the sandy beaches being generally close to glacifluvial deposits. At the NW end of Loch Ryan low cliffs continue round Milleur Point, where the coast is dominated by a 5 m terrace, backed by mounds (kames). Many fissures have been excavated by wave action along the rocky shore. There is a wide coastal ledge, and some segments of higher terraces in bluff re-entrants. Along the shore are a few cobble beaches, some with shell grit, between rocky outcrops and boulders along the seaward edge. South of Dally Bay are rocky headlands and coves (>Fig. 7.24.10.4), then bluffs behind a beach of pink sand, with some low dunes. Knock Bay and Killantringan are set back by a fault. Port Kale and Port Mora are valley-mouth coves with sandy floors, grey gravel beaches, and rocky margins. Portpatrick has rocky shores and a sandy inner and outer harbour in a little inlet. Cliffs extend south past Portayew to the Mull of Logan. Cliffs and rocky shores to the south are variously stratified, gnarled and craggy. The Mull of Galloway is a protruding peninsula on metasediment, with cliffs bordering an isthmus between West Tarbet and East Tarbet bays. The cliffs are higher along the steeper west coast which has greater exposure to storm waves than the eastern side. Drummore has a

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South-West Scotland

sandy beach and segments of rugged rocky shore, backed by hum­mocky drift, few rock outcrops, and there is a sandy beach in the bay at Ardwell. Behind a widening coastal plain are rising slopes in metamorphic rocks (>Fig. 7.24.10.5). At Sandhead the sandy beach and dunes come to an end, and the shingle beach gives place to sand, backed by dunes eastward along Luce Sands (Single and Hansom 1994). The beach is backed by low parallel foredunes, generally advancing seaward, but locally cliffed. Behind these are high transverse dunes and older dunes that rest on emerged shingle beach ridges and show signs of severe erosion in the past (Mather 1979). The sandy beach ends at Stair Haven against cliffs that run south to the Mull of Sinniness. The long east coast of Luce Bay is mainly cliff-bound but there are sandy beaches at Auchenmaig Bay and Craignarget Bay. Cliffs run SE to Burrow Head on an outcrop of metamorphic rocks. To the east the Isle of Whithorn is attached to the mainland by a causeway across what was once a narrow tidal strait. Low cliffs run north to Garlieston Bay, where at low tide wide grey-brown sandy and muddy areas are exposed, with patchy salt marsh. Wigtown Bay narrows to salt marshes beside the muddy estuary of the River Cree (Firth et al. 2000). To the SE the gravelly shore is backed by steepening slopes and then cliffs rising to Ringdoo Point beside Fleet Bay. Fleet Bay has grassy headlands between sandy beach­es and tidal mudflats. Upstream the Water of Fleet estuary has a reedy fringe. On the east coast there are grassy

­ eadlands, as at Sandgreen, and sandy bays which run h behind the Islands of Fleet. The southern shores of Kirkcudbright bay are rugged, and northward become muddy with salt marsh. East of Kirkcudbright Bay cliffs in Middle Silurian rocks run towards Abbey Head, which has a grassy bluff behind a raised beach. Abbey Head can be taken as marking the outer limit of Solway Firth.

4.  Solway Firth East of Abbey Head the cliffs decline to the valley mouth at Abbey Burnfoot, where a stream flows out to the sea through a curving partly vegetated shingle berm behind a bouldery shore. At low tide the sea drains out to expose extensive sand flats backed by shingle beaches in Auchencairn Bay and Fishery Bay, where granite outcrops on the shore. Rough Island is also granitic, and has been attached by an intertidal shingle bar (The Rack) to form a peninsula with Rough Firth on its eastern side. Cliffs rise from Rockcliffe to Castlehill Point, bordered by a sandy beach and pebbly cove. Towards the Elbe monument a scrubby bluff edging a terrace recedes behind an emerged shore platform, the inner half vegetated, with segments of bedding plane dipping eastward. The cliffs have a slope-over-wall profile, with some rocky outcrops but no shore platforms.

⊡⊡ Fig. 7.24.10.5 Cliffs near the Mull of Galloway in Silurian metamorphic rocks. (Courtesy Department of Geography, University of Aberdeen.)

South-West Scotland

7.24.10

⊡⊡ Fig. 7.24.10.6 Pond holes scoured in the salt marsh at Sandyhills Bay. (Courtesy Geostudies.)

East of Port o’Warren is the pebbly cove of Portling Bay. In Sandyhills Bay (>Fig. 7.24.10.6) eroded dunes behind a beach of shelly sand pass into salt marshes which show the typical pattern of an upper marsh with a receding cliff stepping down to a lower marsh, eroded at the seaward edge but with clones of Spartina (Harvey and Allan 1998). On the lower marsh are pond holes with rounded cobbles and boulders which extend beneath the marsh (>Fig. 7.24.10.6). At Southerness the rocky shore has gently dipping Calciferous Sandstone and sandy patches behind the wide intertidal Mersehead Sands. The intertidal Carse sands con­tinue past Carsethorn, a coastal village on a lobate foreland. The shore is fringed by a beach of sand, pebbles and shells. At Burnfoot, on the western shore of the Nith estuary there is a grassy upper marsh with some reedswamp and a darker green lower marsh with an eroded fringe. Tides are strong in the Nith estuary. The large salt marsh that fringes the coast at Caer­ laverock is fronted southward by very wide sand flats exposed at low tide. Terraces have formed as the result of Holocene emergence (Firth et al. 2000).

To the east at Bankend is the mouth of Lochar Water, another marshy estuary. There is a broad gently sloping coastal plain of glacial drift ending seaward in merse (high salt marsh). The shore becomes sandy eastward to Powfoot and towards the mouth of Annan River. On the shore SE of Annan low cliffs of glacial drift have been stabilised behind a narrow fringe of high and low salt marsh. The low lying country east of the Annan estuary is largely reclaimed farmland behind artificial earth em­bankments.

References Firth CR, Collins PEF, Smith DE (2000) Focus on Firths, V – The Solway Firth. Scottish Natural Heritage, Edinburgh, UK Harvey MM, Allan RL (1998) The Solway Firth salt marshes. Scot Geogr Mag 114:42–45 Mather AS (1979) Beaches of Southwest Scotland. Department of Geography, University of Aberdeen, Aberdeen, Scotland Single MB, Hansom JD (1994) Torrs Warren-Luce Sands SSSI: documentation and management prescription. Scottish Natural Heritage, Survey and Monitoring Report, 13, 85 p

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7.25 Northern Ireland

Andrew Cooper

1.  Introduction The coastline of Northern Ireland is dominated by bedrock outcrops of considerable diversity. The southeast coast (County Down) is composed of low relief Palaeozoic shales intruded by the Tertiary Mourne and Newry granites while the northeast and northern Antrim and Londonderry coasts comprise mainly Mesozoic mudstones and chalk with a covering of Tertiary basalt. The Antrim and Lon­ donderry coast has several areas of high relief and impressive cliffs are present at some locations, as at Fairhead. Both rotational slumping of basalt and chalk over underlying clays and periodic rock falls have created impressive coastal slope deposits in association with the cliffed shoreline, while marine erosion has created impressive shore platforms in the basalts. Sand and gravel beaches are located on this bedrock-dominated coast in many small embayments. The largest beach systems are located at Dundrum Bay in the southeast and Magilligan in the northwest. Both are backed by extensive, vegetated sand dunes. The whole of Northern Ireland was affected by tem­ porally variable patterns of ice movement during the Quaternary (McCabe and Dunlop 2006) which, through loading and offloading of the crust, contributed to a complex sea level history during the past 20,000 years. From the perspective of the modern coastal geomorphology, the major influences of the Quaternary ice sheet have been in the deposition of large quantities of sediment on the continental shelf and localised parts of the coast, and in the complex sea level history. A sequence of raised shorelines that formed during the deglaciation that began about 20,000 years ago are preserved at heights up to 30 m above present sea level. At the close of the glacial period about 13,000 years bp, relative sea level was at its lowest position (−30 m). Rapid sea level rise during the early Holocene saw the flooding of the coast until about 7,000 years ago when the sea level exceeded the present by up to 5 m. Since then and into the historical period, sea level has fallen slowly due to ongoing land uplift. Tidal records from Malin Head and Belfast reveal a slowly falling to stable relative sea level (Orford et al. 2006), which indicates a near-balance between the rate of land uplift and the rate of eustatic sea level change.

The fall in sea level over the past 5,000 years has resulted in deposition of large regressive coastal sedimentary sequences at sites where sediment supply has been available to sustain them. In other areas, where the sediment supply has run out, higher than present shorelines have been left stranded above present sea level. The extensive sandy beach ridge plain of Magilligan at the mouth of Lough Foyle was deposited as sea level fell from its late Holocene high stand, while the large dune system at Dundrum Bay was deposited on top of a similar sequence of beach ridges with gravel armouring. In both cases, reworking of former glacial sediment from the continental shelf provided the source of sediment for beach ridge and dune development during the late Holocene. The rock coast of Northern Ireland contains well-­ developed cliff-shore platform units that are particularly  well developed in the basalts of the Antrim coast (McKenna et al. 1992). Here, the sub-horizontal flow tops of the basalt have been preferentially exploited by wave action to create wide shore platforms, the best-known of which is the Giant’s Causeway which is composed of hexagonal basalt blocks and columns. Associated with the shore platforms are boulder beaches that have developed through accumulation of eroded basalt blocks in coastal re-entrants. Other notable features of the solid rock coast of Northern Ireland are the rotational slumps of chalk and l basalt that are common along the Antrim coast and the erosional features (stacks, arches, caves) that are found in the chalk and basalt. For its size, Northern Ireland exhibits a great diversity of contemporary dynamics (Jackson et al. 2005; Cooper 2006). It is surrounded by a shallow shelf sea influenced by the North Atlantic in the west and the Irish Sea in the east. Strong tidal currents are present on the shelf. Spring tidal range is >2.5 m at Lough Foyle, falls to Fig. 7.25.1). These sand beaches have well-developed ridge and runnel morphology comprising up to five intertidal to subtidal ridges across a 500-m wide beach. A gravel storm beach is present along much of the high tide shoreline. This is succeeded landward by a 1-km-wide system of high and irregular sand dunes. The entire system is bisected by a tidal inlet that connects Dundrum inner bay to the sea. It has a well-developed ebb-tidal delta at its seaward limit. On the western (Newcastle) side of the tidal inlet the intertidal shore commonly exhibits up to six low parallel ridges separated by intervening troughs, forming a classic ridge and runnel topography. The beach here exists as a veneer of sand overlying a compact underlying surface topped by cobbles. The underlying cobbles are well exposed in the vicinity of Newcastle. The intertidal beach on the eastern (Ballykinler) side of the bay is a wide, planar feature backed by a series of shore-parallel foredunes and an older system of high hummocky dunes. A series of wide bedrock rock platforms front low, generally inactive bluffs cut into glacial sediment around St. Johns Point. To the west is a wide shore platform with occasional patches of sandy beach trapped in small embayments. The extent of intertidal sand gradually increases towards Tyrella, where a wide beach is backed by extensive dunes. The stretch of coast between Ringfad Point and Ballyhornan is dominated by bedrock outcrops and is almost devoid of modern sediment. Much of the coast contains a raised beach perched on a probably wave-planed shore platform above present sea level. Ardglass is a rocky embayment that has a few gravel accumulations in gullies around its margins, but little mobile sediment is present. At Ballyhornan two small sandy beaches are separated by a rocky headland composed of wave-eroded Palaeozoic

Northern Ireland

7. 25

⊡⊡ Fig. 7.25.1 Intertidal sand beach backed by gravel storm beach in Dundrum Bay. The Tertiary Mourne Granites form Northern Ireland’s highest mountain (Slieve Donard) in the background.

bedrock on which rests a raised gravel beach. The modern beaches are thin veneers of sand and subordinate gravel that overlie a wave-eroded glacial surface. At the rear of the beach is a slowly eroding cliff cut into glacial deposits from which eroded blocks have fallen and formed boulder accumulations. The short stretch of coast between Dundrum Bay and the mouth of Strangford Lough is predominantly rocky. The only stretches of sandy coast are within an embayment at Killough and at Ballyhornan. Elsewhere, the near-linear coast offers little accommodation space for accumulation of any available sediment. The absence of sandy deposits in Ardglass harbour also points to a scarcity of sediment supply along this coast. Raised beach deposits are more common than modern beaches, suggesting that much of the available sediment is now stranded above modern wave processes as a result of late Holocene sea level fall. Strangford Lough is a narrow marine embayment into which small rivers discharge. Its most prominent features are the drumlins that have been isolated by rising sea levels to form numerous small islands of glacial material. Erosion of the drumlins by contemporary wave action is comparatively slow (Greenwood and Orford 2007) but in some instances, gravel ridges, composed of coarse clasts eroded from drumlins have accumulated. At various ­locations around the lough small areas of salt marsh and tidal flats exist, although much of the former extent of these features has been lost through land reclamation.

The greatest extent of tidal flats is in the northern section of the Lough. In general, the more exposed tidal flats on the northern and eastern shores of the lough are sandy while those in sheltered, typically eastern locations are muddy. Strong tidal currents through the rock-bounded channel that connects the lough to the sea are associated with its large tidal prism. At Cloghy a wide sandy beach is located between headlands at Slanes Point in the south and Ringboy Point in the north. The intertidal sand beach is 200 m wide at  low tide with a narrow (10–15 m wide) gravel storm beach at high tide. Virtually the whole beach is backed by vegetated dunes. “The Ridge” which extends seaward of Ringboy Point is a kilometre-long deposit of coarse gravel and boulders that is probably the reworked upper surface of a glacial moraine. It provides an obstacle to northward transport of sand and has assisted the trapping of sediment in the bay. South of Cloghy is a long stretch of rocky coast indicating that there is little potential additional sand available for continued beach growth. At Millisle beaches of grey sand are interspersed with outcrops of bedrock. An extensive sand beach south of the town has formed largely in the shelter of Long Isle, an elongate offshore rock outcrop that runs parallel to the coast and provides a natural offshore breakwater. The reduction in wave energy has enabled sand to accumulate in its lee. The upper beach is 20–30 m wide but is fronted by an intertidal flat 100 m wide.

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Northern Ireland

The Ards Peninsula coastline consist of alternations of low-relief Palaeozoic shale outcrops with gravel and sand beaches of variable thickness and lateral extent. Longshore drift cells were demonstrated by Bowden and Orford (1984). The most impressive sandy beaches on this stretch are at Millisle and Cloghy. The main A2 road parallels the coast for much of the length of the outer Ards peninsula and where it impinges on beaches it has invariably been armoured. Ballyholme near Bangor has a sandy beach located between two prominent headlands, backed by a seawall and promenade. Holocene wetland peats are exposed under the sand on the modern beach which forms a thin veneer. The outer coastline of north County Down and the Ards peninsula is dominated by relatively low-lying rock outcrops consisting mainly of folded and faulted Palaeo­ zoic mudstones and sandstones. Much of this bedrock is covered by Pleistocene glacial deposits, often in the form of drumlins. The Palaeozoic bedrock is exposed where the overlying sediment have been removed by wave action. Erosion of the Pleistocene sediment onshore and on the seabed has provided the source material for beach building on this coast. A veneer of sandy or gravelly sediment is present along much of the coast, the more substantial beaches occuring in indented embayments and in the lee of offshore rock outcrops. Raised beaches of postglacial age are found intermittently along the outer Ards coast and in some cases they provide a sediment source for modern beaches through coastal erosion. The fall in sea level from a postglacial high stand a few metres above present has stranded formerly active eroding bluffs of glacial material.

3. The Northeast Coast (Belfast Lough to Ballycastle) The coast curves westward to Bangor and into Belfast Lough, a shallow glaciated embayment that was drowned during the Holocene rise in sea level. It has been extensively modified by land reclamation and port development along its upper reaches. The bed of the lough is dominated by muddy sediment which blanket shoreline and estuarine deposits from lower Holocene sea levels. Small sandy beaches are present around the Lough, particularly on its southern shoreline at Holywood and Crawfordsburn where a small stream flows across the beach. To the north the Gobbins cliffs consist of steep basalt into which are cut a series of caves and notches that represent marine erosion at higher sea levels. A discontinuous shore platform of variable width is present along the base of the cliffs.

Larne Lough is separated from the open ocean by the peninsula of Islandmagee. The Lough contains extensive intertidal mudflats and saltmarsh is well developed in its upper reaches. The coastline of Islandmagee consists mainly of basalt cliffs with and local chalk outcrops. These have been wave eroded and in several locations, extensive abrasion platforms, notches and caves have been cut by wave action at a variety of higher sea levels during and since deglaciation. Eroded basalt blocks forms local accumulations in embayments and on raised platforms around the coast. A raised Late Holocene shoreline is particularly well developed around the northern end of Islandmagee. Sand and gravel beaches are present in the larger embayments on this otherwise rocky coast at Brown’s Bay, Portmuck and White­ head. Portmuck (>Fig. 7.25.2) is a small embayment with a narrow upper pebble beach of chalk and flint with a few basalt pebbles, fronted by a fine sandy tidal flat. The upper beach is partly vegetated, and backed by a raised beach. The modern beach sediment is derived mainly from erosion of the local chalk. Between the Islandmagee and the Isle of Muck is a linking bar covered even at low tide in its central portion but the extremities are above high tide level. It has a core of basalt, chalk and flint boulders, around 50 cm in diameter, and is thinly covered on its north eastern side by smaller pebbles (5 cm in diameter). Its south-facing margin comprises coarser clasts. It is steeper on its concave north east face. Most of the pebbles are chalk and to a lesser extent basalt, there are some exotic lithologies (sandstones, porphyry and granite) which are probably derived from glacial deposits. At Larne, the largest accumulation of beach gravels on the northeast coast is present at Curran Point. This gravel spit has been heavily modified by construction associated with Larne Port (Carter 1991). Between Larne and Ballycastle is a coastline of moderate relief cut by a series of valleys (the Antrim Glens). The coast is predominantly rocky with alternations of basalt, chalk and mudstone. In the embayments that are mainly associated with valley entrances, small beach systems are developed. These are of varying texture, ranging from well-sorted sand to medium gravel with various mixtures of  these constituents. The main beaches are present at Ballygalley, Glenarm, Red Bay, Cushendall and Cush­ endun. Each is backed by a small vegetated dune field. Each of these beaches has been affected to some extent by human activity because of the coincidence of ­villages with valley entrances and the presence of the Antrim Coast Road (Carter 1991). Although it is widely regarded as a great feat of engineering and a scenic drive, this road

Northern Ireland

7. 25

⊡⊡ Fig. 7.25.2 Chalk coast at Portmuck Bay, Islandmagee. (Courtesy Geostudies.)

⊡⊡ Fig. 7.25.3 Chalk shore at Garron Point. (Courtesy Geostudies.)

has ­interrupted the transfer of sediment along and across the shore. Em­­bankments at the rear of several beaches have ­separated beaches and dunes and caused beach lowering; they are damaged during storms and require frequent maintenance. Elsewhere the cliffs that formerly supplied sediment to the coastal system are now physically separated from it. Landslides occasionally cause the road to be closed. The rock coast of northeast Antrim contains a wealth of slope forms and deposits. Mudflows (Prior et al. 1968)

are found in the soft lithologies of Jurassic and Triassic age and are particularly well developed at Murlough Bay. Large scale rotational slumping is also a characteristic of this coast. The more competent chalk and basalt slide over the underlying Jurassic clays, and slumped landward-dipping chalk and basalt and are common north of Garron Point (>Figs. 7.25.3 and > 7.25.4). Raised coastal landforms of late Holocene age are common along the northeast Antrim coast. These comprise both depositional forms (beaches and barriers) and

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Northern Ireland

⊡⊡ Fig. 7.25.4 Dissected Chalk shore ramp at Garron Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.25.5 Basalt slopes at Torr Head. (Courtesy Geostudies.)

erosional forms (stacks and caves) as at Cushendun on a localised outcrop of Devonian conglomerate. The basalt coast extends past Torr Head (>Fig. 7.25.5).

4. The North Coast (Ballycastle to Magilligan) The cliffs at Fair Head (>Fig. 7.25.6) just east of Ballycastle mark the divide between the north-eastern coast on which

waves generated by winds within the Irish Sea predominate and the north coast where North Atlantic Swell waves are dominant. Offshore Rathlin Island has impressive basalt and chalk cliffs. The rocky north coast contains some active and relict high cliffs cut in the basalt and chalk outcrops and other erosional landforms. Shore platforms are well developed in the basalt where storm waves have exploited the subhorizontal surfaces between successive lava flows. Potholes and irregular-shaped pools are common where joints and

Northern Ireland

7. 25

⊡⊡ Fig. 7.25.6 Basaltic cliffs and scree slopes at Fair Head, Ballycastle.

⊡⊡ Fig. 7.25.7 Coastal terrace at Ballintoy. (Courtesy Geostudies.)

other weaknesses are exploited by wave action. Rework­­ ing  of eroded cliff and platform boulders results in the accumulation of rounded boulder beaches in adjacent ­embayments. By contrast coastal outcrops of white chalk have more limited shore platforms, but caves, arches and stacks are more common. A number of small rocky islands are present along the north coast, the best known of which is Carrick-a-Rede, west of Ballycastle, which is seasonally accessible by a rope bridge. There is a coastal terrace and outlying basaltic islands at Ballintoy (>Fig. 7.25.7). The north coast between Ballycastle and Portrush ­consists of a series of rocky cliffs interspersed with

­ redominantly sandy beaches and sand dunes in coastal p re-entrants at Ballycastle, Whitepark Bay, Portballintrae and Portrush. Each of these beaches is confined between rocky headlands  and is composed of sand derived from the continental shelf during the early Holocene rise in sea level and reworked during the past 7,000 years of slightly falling to stable sea level. Extensive dunes are developed at all of these beaches. During storms dune erosion may occur, but the confined nature of these beaches and their finite sediment volume means that in post-storm periods the eroded material is returned to the beach. Erosion problems do occur, ­however, as a result of human ­activities including

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Northern Ireland

⊡⊡ Fig. 7.25.8 Chalk cliffs penetrated by basalt east of Portrush. (Courtesy Geostudies.)

⊡⊡ Fig. 7.25.9 Caves excavated along joint planes in Chalk near Portrush. (Courtesy Geostudies.)

sand removal from beaches and engineering works. At Portballintrae a ­formerly wide sandy beach has been eroded as a result of pier construction, and despite several attempts to reinstate the beach (groynes and small scale nourishment) it remains a poor remnant. The Chalk cliffs at Portrush are capped by basalt, which locally penetrates down into them (>Fig. 7.25.8), and are dissected by caves (>Fig. 7.25.9). At Portrush West Strand, lowering of the beach after storms often reveals a freshwater peat outcrop under the beach (>Fig.  7.25.10). This beach has become narrower

and lower since a seawall built during the 1960s separated the beach from the dunes. West of Portrush the coast curves out to the Magilligan Foreland. Wide sandy beaches at Portstewart and Cast­ lerock are separated from similar beaches at Downhill and Benone by a short section of basalt cliff. All these beaches are backed by extensive vegetated sand dunes and cliff-top dunes are also present on adjacent rocky coasts. The ­estuary of the River Bann separates Portstewart and Castlerock beaches and its inlet has been stabilised by the

Northern Ireland

7. 25

⊡⊡ Fig. 7.25.10 The sandy beach at Portrush West Strand has narrowed and lowered since the seawall and promenade were constructed. The peat underlying the beach is exposed periodically after storms.

⊡⊡ Fig. 7.25.11 Fresh foredune dune accumulation at Magilligan Point. This coastline has long periods of erosion followed by accretion.

construction of training walls in the late nineteenth century. The high, vegetated dunes around the estuary are among the largest in Northern Ireland and they contain a record of environmental changes over the past 7,000 years (Wilson and McKenna 1996).

Magilligan Foreland (>Fig.  7.25.11) is composed of a series of sandy beach ridges (Carter 1986). To the south is Lough Foyle, a large, shallow sandy estuary, and to the west a 1,600-m-wide tidal inlet through which the estuary ex­­ changes water with the North Atlantic. A large ebb-tidal

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Northern Ireland

delta (the Tunns Bank) is present offshore and Carter et al. (1982) described the cyclic exchange of sediment between the tip of Magilligan Point and the ebb delta. When the delta has maximum amounts of sand its crest becomes emergent, while the adjacent shoreline at Magilligan Point retreats.

References Bowden R, Orford JD (1984) Residual sediment cells on the morphologically irregular coastline of the Ards Peninsula, Northern Ireland, Proc Roy Irish Acad 84B:13–27 Carter RWG (1986) The morphodynamics of beach-ridge formation: Magilligan, Northern Ireland. Mar Geol 73:191–214  Carter RWG (1991) Shifting Sands: a study of the coast of Northern Ireland from Magilligan to Larne. HMSO, Belfast Carter RWG, Lowry P, Stone G (1982) Ebb shoal control of shoreline ­erosion via wave refraction, Magilligan Point, Northern Ireland. Mar Geol 48:M17–M25

Cooper JAG (2006) Geomorphology of Irish Estuaries: inherited and dynamic controls. J Coastal Res Special Issue 39:176–180 Cooper JAG, Navas F (2004) Natural bathymetric change as a control on century-scale shoreline behavior. Geology 32:513–516 Greenwood RO, Orford JD (2007) Factors controlling the retreat of drumlin coastal cliffs in a low energy marine environment-Strangford Lough, Northern Ireland. J Coastal Res 23:285–297 Jackson DWT, Cooper JAG, Del Rio L (2005) Geological control on beach state. Mar Geol 216:297–314 McCabe AM, Dunlop (2006) The last glacial termination in Northern Ireland. Geological Survey of Northern Ireland, Belfast McKenna J, Carter RWG, Bartlett D (1992) Coast erosion in north-east Ireland: Part II, cliffs and shore platforms. Ir Geog 25:111–128 Orford JD, Murdy J, Freel R (2006) Developing constraints on the relative sea-level curve for the north-east of Ireland from the mid-Holocene to the present day. Phil Trans Roy Soc A 364:857–866 Prior DB, Stephens N, Archer DR (1968) Composite mudflows on the Antrim coast of north-east Ireland. Geogr Ann 50A:65–78 Shaw J (1985) Beach morphodynamics of an Atlantic coastal embayment: Runkerry Strand, County Antrim. Ir Geogr 18:51–58 Wilson P, McKenna J (1996) Holocene evolution of the River Bann ­estuary and adjacent coast, Northern Ireland. Proc Geol Assoc 107:241–252

7.26  Ireland

1.  Introduction The Republic of Ireland has an area of 70,282 sq. km, and consists of a large Central Lowland (generally less than 120 m above sea level) bordered by a hilly and mountainous coastal fringe. The coastline is controlled by the influence of several major geological structures, notably the NE-SW Caledonian structures in the northwest of the island and the E-W Armorican structures in the south. Mountain ranges and deep valleys following these trends culminate in long peninsulas and bays on the Atlantic coast in the northwest and southwest, the limestones of the interior extending to the central coast in Galway. The south coast runs parallel to the Armorican trend in County Cork, and the southeast and east coasts are hilly with some valley-mouth lowlands (Davies and Stephens 1978). In the W and NW are Pre-Cambrian schists and gneisses similar to those of the Scottish Highlands; in the NE are Ordovician and Silurian rocks like those of the Southern Uplands of Scotland and in the SE Ordovician rocks equivalent to those of western Wales. The major intruded granites include the Leinster granite which reaches the south coast of Dublin Bay, the Galway granite which runs along the north coast of Galway Bay and the Donegal granite along the coast south from Bloody Foreland. Old Red Sandstone (Devonian) occurs in the SW along the long narrow Dingle, Inveragh and Beara Peninsulas, and in the south in anticlinal ridges separating synclinal vales of Carboniferous Limestone east and west of Cork. Carboniferous limestones, shales and sandstones dominate much of central Ireland, extending to the coast at and north of Dublin Bay, between Donegal Bay and Killala Bay, in Clew Bay, and at and south of Galway Bay. They also form The Burren south of the Shannon estuary, and outcrop in Tralee Bay and Dingle Bay, in short sectors in Ballycotton Bay, Youghal Bay and Dungarvan Harbour on the south coast, and in the south-east at Ballyteige Bay and Wexford Harbour. Much of Ireland bears the imprint of Pleistocene glaciation, both in the eroded uplands and broad

drift-mantled lowlands, and also periglaciation. Ice spread over the whole of Ireland in the Munsterian (Wolstonian) glaciation, and over much of the country in the Midlandian (Devensian) glaciation, when the coast between SW Clare and the western end of the Iveragh Peninsula remained ice-free, as did the south coast east of Ballycotton Bay. Many coastal features have been shaped by glacial and periglacial processes, including slope-over-wall cliff profiles, fiords and fiards, cliffs cut in glacial drift deposits and beaches derived from moraines, drumlins and eskers. Coastal features have also been shaped by as well as sea level alternations culminating in the Late Quaternary (Flandrian) marine transgression and postglacial isostatic uplift, notably in the northwest and east of the country. The Atlantic coasts are rugged and rocky where the glacial drift cover has been removed during the 6,000 years since the last marine transgression rise brought the sea to about its present level. Some are relict cliffs that originally formed in Pleistocene times, and have been exhumed from a drift cover. By contrast, much of the east coast remains drift-mantled. The east coast is generally low, with few headlands exceeding 30 m in height. Raised beaches, indicating higher relative sea levels in Quaternary times, decline in level southward as the result of postglacial isostatic rebound: the 7.5 m raised beach in SW Scotland declines to 3 m near Dublin and intersects present sea level near Arklow. The older and higher (Late Glacial) raised beaches show evidence of frost heaving, contain no shells and often rest on benches cut into glacial drift, in each of these respect contrasting with the younger and lower (Postglacial) raised beaches. There are submerged forests, notably at Bray and Wicklow (Stephens 1970). Land has been claimed from the sea on parts of the coast, particularly in sheltered bays and inlets, but erosion has been taking place on many of the beaches, especially where there has been removal of sand, shells and seaweed. There has been extensive construction of sea walls and boulder ramparts.

Edited version of a chapter by N. Stephens in The World’s Coastline (1985: 377–383). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.26, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The climate of Ireland is mild and maritime, with mean annual rainfall up to 3,000 mm on the west coast, diminishing to 780 mm in Dublin on the east coast, where the mean monthly temperature increases from 4.7°C in January to 15°C in July. Peat deposits are extensive in the wetter areas, reaching the coast locally in Donegal. On the south coast Cork has a mean monthly temperature of 6°C in January and 15.6°C in July and annual rainfall of 1,025 mm. Mean spring tide range is generally between 3 and 4 m, rising to 4.5 m at Galway and Kilrush on the west coast. On the east coast Wexford Harbour has 1.5 m, Dublin 3.6  m and Dundalk 4.7 m. In the north-west Inistrahull has 2.9 m and Sligo Harbour 3.6 m. On the west coast Broad Haven has 3.2 m, Clare Island 3.6 m, Inishbofin 3.6 m, Galway 4.5 m, Kilrush 4.5 m, Tarbert Island 4.5 m, Tralee Bay 3.9 m, Dingle Harbour 3.3 m, Ballinskelligs Bay 3.1 m and Bantry 2.9 m. On the south coast Castle Townsend has 3.3 m, Cobh 3.7 m, Ballycotton 3.6 m, Youghal 3.6 m, Dungarvan 3.7 m and Dunmore 3.6 m.

2.  The Coastline of Ireland Lough Foyle, in NW Ireland, is a large marine inlet occupying a synclinal trough. It has a tide range of about 2 m, but wave energy is low because the entrance is constricted by the large Magilligan Foreland to the north. There are several delta-like lobes at stream mouths along the coast, as at Quigley’s Point, but although they have the ­appearance

of deltas on maps and air photographs they are actually residual lobes on a receding coast where low cliffs have been cut in an apron terrace of glacial drift. The lobes persist at stream mouths because the adjacent nearshore areas have been shallowed by outflow deposition (essentially intertidal deltas), thus impeding wave attack, which has cut back intervening sectors further. On the coast of Donegal the NE-SW Caledonian trend exerts a strong geological control on the coastline. Steep coast profiles have been shaped partly by Pleistocene glacial and periglacial action, with basal cliffs cut by wave erosion in the 6,000 years since the Late Quaternary (Flandrian) marine transgression brought the sea to roughly its present level. Inishowen Head shows a coastal slope on glacial and periglacial drift descending to a basal cliff cut in schists and quartzite and similar profiles are seen on Ballane Head, west of which inwashed calcareous sand forms beaches and dunes. There are remnants of Late Glacial and Postglacial emerged beaches. Intensive scouring and excavation by glaciers has shaped mountains and valleys, and the Late Quaternary marine transgression has produced the partially drowned and island-obstructed fiards at Sheep Haven, Mulroy Bay and Lough Swilly, between Gweebarra Bay and Malin Head. Around the mouths of each of these rias are bold, high cliffs that diminish inward, and sand-fringed shores passing upstream to mudflats and salt marshes. Some bordering bays have been reclaimed as farmland, but north of Rathmelton one has been abandoned. Shelly and calcareous sand beaches fringe the west coast of Lough Swilly at ⊡⊡ Fig. 7.26.1 The coast east of Malin Head, Donegal, showing emerged gravelly marine terraces. A scarp marks a former cliff line behind the lowest terrace, which is of Holocene (postglacial) age. The wide gravel beach consists of cobbles and pebbles of a variety of rock types, derived from glacial drift. (Courtesy Geostudies.)

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Rathmullan and in Ballymacstocker Bay. Sandy beaches also occupy valley-mouth bays, as at Ballyniernan. In Sheep Haven calcareous sandy beaches are backed by dunes, which have been truncated by marine erosion at Five Finger Strand. Late Glacial, Postglacial and modern storm beaches of pebbles and cobbles occur at heights of up to 25 m on Malin Head (>Fig. 7.26.1), Fanad Head, Horn Head and the Bloody Foreland. Malin Head is a promontory of hard quartzite with storm-piled shingle beaches derived from glacial drift, and similar features are seen on Fanad Head and Horn Head, whereas Bloody Foreland is granitic, with slopes descending to low cliffs. Stacks rise from the rocky shore platform at Five Finger Strand. There are beaches composed mainly of quartz and shell sand, which are particularly well developed at Pollan Bay, Ballyheirnan Bay and the barrier spits at Ballyness. Those occupying bays on the outer coast are curving strands shaped by refracted northwesterly swell from the Atlantic (Duffy and Devoy 1999). Blown sand has proved to be a menace to farmland and settlements at Dunfanaghy and Rosapenna, southeast of Horn Head, where sand blowing has been a serious problem at least since 1784. At Rosapenna a sheet of windblown sand up to 6 m thick overwhelmed numerous farms after the coastal dunes were destabilised by rabbit grazing following control of foxes. Sand has been washed into branches of Sheep Haven, as at the Back Strand. On the Dooey Peninsula a major cuspate barrier, capped by dunes, is interrupted by the tidal entrance to Ballyness Bay, a lagoon at high tide and an area of wide ⊡⊡ Fig. 7.26.2 The megacliff at Slieve League, Donegal. (Courtesy Geostudies.)

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intertidal sand and mudflats at low tide. The barrier runs out as a sandy spit shaped by convergent waves in the lee of Inishbofin Island. Offshore to the north is Tory Island, granitic, with a quartzite tip at its eastern end. The land rises towards the north-east coast of this island (83 m) where a chasm (geo) runs northward beside a narrow peninsula at The Anvil. The Bloody Foreland is a cliffed granitic promontory with emerged Late Pleistocene and Holocene gravel beaches extending more than 20 m above sea level. To the west, Innishfree Bay has sandy beaches, massive gravelly storm beaches and spits derived from glacial drift, illustrating sorting by waves and currents as longshore drifting proceeds. There are sandy tombolos, one of which carries Donegal Airport. In western Donegal sand has been washed into the mouths of long valley-mouth inlets to form intertidal thresh­olds that pass to muddy sediment upstream, towards the mouths of inflowing rivers. Maghera Strand, south of the Loughros Peninsula, consists of pale calcareous sand that has been swept in from the sea floor by ocean swell. The upland peninsula in southwest Donegal has sectors of high, steep coast, as at Slieve Tooey and Glen Head. Steep coastal slopes show evidence of instability in the form of slumping, subsiding terracettes and gully erosion. Beaches in bays and coves are dominated by calcareous sand from the sea floor, often with coarser gravel derived from glacial drift. On Slieve League (598 m) in southwest Donegal steep (30° to 60°) scree-clad slopes descend to vertical

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cliffs 100–200 m high (>Fig. 7.26.2). It is a megacliff of quartzite with a capping of metamorphosed limestone (Guilcher 1966). The profile has been partly shaped by glaciation, for the upper slopes are similar to the ice-plucked corrie wall of Lough Agh to the north, but quartzite buttress tors up to 12 m high rise above the scree in places and constitute part of the late Pleistocene periglacial modification of the coastal slope. Only the lower cliff is of marine origin, cut back by storm wave action. In southern Donegal long peninsulas of Carbonif­­ erous  sandstone and limestone show cliffs bordered by structural shore platforms, as at St. Johns Point and Doorin Point. At the head of Donegal Bay drumlins composed of  glacial till have been partly submerged to form ovoid hilly islands, and cliffs cut into these have yielded sand and gravel, deposited as beaches and spits, some capped by dunes. There are some fragments of Postglacial raised beaches 1–3 m above mean sea level, as at Eossnowlagh Lower. South of Donegal cliffed drumlins form headlands  north and south of the wide sandy Murvagh Beach, which is backed by a broad dune area. The gentle beach profile shows high tide strandlines and swash bars (>Fig. 7.26.3). South of Bundoran in Leitrim Middle Carboniferous limestones, sandstones and shales outcrop along the coast in low cliffs and shore platforms. On the Mullaghmore promontory the cliffs and platforms show dissection of sandstones by wave quarrying and corrosion features (pits and channels) on limestones. At Grange a tidal lagoon has been almost enclosed by a dune-capped barrier spit. The

wide drift-mantled coastal plain extends back to the Benbulbin escarpment, and beaches and spits derived from glacial drift border the shores of Sligo Bay. Emerged shore platforms and beaches are evidence of former higher sea levels, as at Culleenamore, where a beach overlies a kitchen midden with oyster shells. To the west the northfacing coast of Sligo is generally low-lying, fringed by low cliffs and shore platforms on Carboniferous Lime-stone. On the eastern coast of Killala Bay from Lenadoon Point south to Inishcrone 35 dykes of early Tertiary basalt run across the limestone shore, and storm waves have piled up a beach of limestone cobbles and riven it landward. In the south of this bay a dune-capped barrier island (Bartragh Island) shelters tidal mudflats in the estuary of the River Moy. To the west, in Mayo, vertical and overhanging cliffs of Carboniferous sandstone rise to Downpatrick Head, a promontory where the stratified rocks have been cut back along joint planes to form clefts, caves, blowholes and stacks (>Fig. 7.26.4). West of Beldberg the sandstones give place to Pre-Cambrian schists, gneisses, quartzites and granites in cliffs up to 260 m high along the coast to Benwee Head. The cliff crest is backed by a southward slope strewn with glacial and periglacial drift derived from higher lost land to the north. There are joint-guided clefts at Porturlin and Moista Sound is a deep chasm (geo) cut into the cliffs by marine erosion along a weathered dyke: it is 120 m deep, 180 m long and only a few metres wide. Benwee Head has a prominent natural arch, and there are several large stacks including Pig Island, Doonvinalla and ⊡⊡ Fig. 7.26.3 Murvagh Beach, south of Donegal, is wide and sandy, backed by grassy and shrubby dunes, with a cliffed drumlin in the distance. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.26.4 The cliff at Downpatrick Head, Mayo, cut in layers of Carboniferous sandstone and shale. (Courtesy Geostudies.)

⊡⊡ Fig. 7.26.5 Cliffs cut in peat on the shores of Tullaghan Bay, bordering Blacksod Bay, where the blanket bogs of County Mayo descend to the coast. (Courtesy Geostudies.)

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the outlying Stags of Broad Haven, which are five small steep-sided high islands, the highest rising to 97 m. To the south at Ross Port beaches border a threshold of inwashed sand that constricts the mouth of an elongated bay. West of Benwee Head Broad Haven separates the Mullet Peninsula, where gneisses and schists form rugged cliffs and dunes have spilled in from west coast beaches. The intermittency of dune movement is indicated by buried soil horizons, and there are areas of low, gently undulating calcareous sand similar to the machair of the Hebrides. The southern tip is of granite. Broad Haven is bordered by an intricate low cliffed coastline. The Mullet Peninsula is attached to the mainland by a narrow isthmus at Belmullet, south of which is Blacksod Bay, where (as in Broad Sound) the blanket bogs of Mayo come down to the sea, and have been cut back as peat cliffs up to 2 m high (>Fig. 7.26.5). Achill Island has steep coastal slopes in Pre-Cambrian quartzite and schist, as at Croaghaun with cirques formed by glacial action. To the east of Croaghaun is Nakeeroge, a corrie lake separated from the sea by a moraine which has been narrowed by marine cliffing, and will in due course be breached to form a marine bay. Clare Island to the south has slope-over-wall cliffs over 300 m high. Clew Bay (>Fig. 7.26.6) is a downfaulted trough partly occupied by an archipelago of drumlins, some with cliffs 10–30 m high and some connected by complex shingle spits formed of gravel derived from the glacial drift. To the south submergence of glaciated valleys has formed long

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⊡⊡ Fig. 7.26.6 The drumlin islands of Clew Bay, Mayo. (Courtesy Geostudies.)

inlets such as Killary Harbour, at the heads of which are intertidal mudflats and salt marshes. On Rinvyle Point an abrasion shore platform is being exposed as overlying glacial drift deposits are cut back in receding cliffs. The intricate coastline of Connemara has numerous headlands and islands of schist and gneiss, and inlets with spits and tidal marshes (Guilcher and King 1961). The north coast of Galway Bay is on granite with numerous ice eroded roches moutonnées, and towards Roundstone the coastline becomes deeply indented and island studded, with outlines determined by joint patterns. Shore platforms are poorly developed, and there are isolated sand and shell beaches dominated by foraminifera, backed locally by low dunes. At the head of Galway Bay Carboniferous Limestone comes to the coast, with a patchy mantle of glacial drift, and there are scattered low postglacial emerged beaches. Kinvarra Bay and Aughinish Bay are inlets fringed by dunes and glacial drift, with some enclosed lagoons as at Finavarra. To the west the south coast of Galway Bay is dominated by the gently tilted Carboniferous Limestone of the Burren and the outlying Aran Islands, which have vertical cliffs up to 90 m high, rock falls and blowholes guided by vertical jointing. Limestone gravels have been spread along the coastline and shaped into beaches. The Burren coast has stepped limestone cliffs with structural ledges, and at Black Head steep limestone slopes descend to karstic coastal ledges and beaches of limestone gravel (>Fig. 7.26.7). Shore outcrops of limestone have been

i­ ntricately weathered by solution processes, and there are elevated sea caves. South of the Aille valley the Burren limestones pass  beneath Carboniferous sandstones and shales form an  8  km sector of high (up to 200 m) cliffs at Moher (>Fig. 7.26.8), their outlines related to joints and beddingplanes (Guilcher 1966). Cliff-foot talus has been worked by wave action into gravelly beaches and there are no shore platforms. To the south the cliffs are lower, and there are bay beaches derived largely from sandy outwash at the limits of the Last Glaciation, which intersect the coast near Spanish Point. Sandstone cliffs and shore platforms continue south-west to Loop Head, with natural arches at the Bridges of Ross. The River Shannon, which has a large catchment in the Central Lowland, drains into a long widening estuary below Limerick. It follows the geological strike, and is bordered by low cliffs cut in glacial drift. At Scattery Island the end moraine of the Midland Glaciation crosses the Shannon estuary. To the south the Upper Carboniferous sandstones and shales of the Ballybunion coast give place to the Old Red Sandstone cliffs of Kerry Head. Carboniferous Limestone returns to the coast in Tralee Bay, where there are major spits at Castlegregory. A mountainous peninsula of Old Red Sandstone runs out to Slea Head and Great Blasket Island, and Dingle Bay is the first of a series of major rias separated by anticlinal peninsulas of Old Red Sandstone: Dingle Bay, Kenmare River, Bantry Bay, Dunmanus Bay and Roaring Water Bay. Towards the head

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⊡⊡ Fig. 7.26.7 Limestone gravel borders a coastal terrace on the Burren coast south of Black Head. (Courtesy Geostudies.)

⊡⊡ Fig. 7.26.8 The megacliffs of Moher, County Clare. (Courtesy Geostudies.)

of Dingle Bay extensive sand deposition has formed intertidal flats and there are paired dune-capped spits at Inch and Rossbeigh (>Fig. 7.26.9). Drumlins are prominent at the head of Bantry Bay. The rias have varying amounts of alluvial fill, while the  intervening peninsulas are flanked by steep slopes and  basal cliffs which grow larger as exposure to storm waves increases southwestward, culminating in high rocky ­headlands, as at Mizen Head (>Fig. 7.26.10). These rias

and peninsulas follow the (E-W) Armorican trend in the rock formations south of the thrust zone which extends from the southern side of Dingle Bay across to Dungarvan on the south coast. Barley Cove, east of Mizen Head, has a sandy beach backed by dunes and low herbaceous plains. Cape Clear marks the beginning of the south coast of Ireland, which runs generally parallel to the Armorican structures. At Baltimore clefts have been cut into the cliffs

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⊡⊡ Fig. 7.26.9 The recurved spit at Rossbeigh, Dingle Bay, SW Ireland. (Courtesy Geostudies.)

⊡⊡ Fig. 7.26.10 Wave-worn outcrops of strongly folded Old Red Sandstone on Mizzen Head, SW Ireland. (Courtesy Geostudies.)

along the strike of strongly folded Old Red Sandstone, while valleys have been incised across the strike. Lough Hyne occupies an ice-scoured valley, and is about 50 m deep, linked to the sea by a narrow and shallow channel through a rocky ridge, where the tides ebb and flow in rapids (>Fig. 7.26.11). The south coast of County Cork has sloping (bevelled) cliffs rising to a broad dissected plateau 80–120 m above

sea level. The coastal slopes are generally in glacial or periglacial drift, descending to low cliffs and shore platforms, mainly in Carboniferous sandstones. There is evidence of Pleistocene emerged coastline features in several places. At Howe Strand, Courtmacsherry, the sandstone cliffs have been cut back to expose an emerged Pleistocene interglacial beach 4–8 m above low tide level, overlain by drift deposits and resting on a partly exhumed

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⊡⊡ Fig. 7.26.11 Rapids produced in the narrow outflow channel from Lough Hyne at low tide. (Courtesy Geostudies.)

⊡⊡ Fig. 7.26.12 The emerged Pleistocene shore platform cut across dipping Carboniferous sandstone of the coast at Howe Strand, Courtmacsherry, west County Cork, is backed by a bluff cut in glacial drift over emerged beach deposits. (Courtesy Geostudies.)

shore platform (>Fig. 7.26.12). The drift deposits include two periglacial Head deposits and an intervening glacial boulder clay resting on the emerged beach. There are reports of a 1930 tsunami at Tragumna, when waves swept up to 2 km inland. There are several prominent headlands, such as Old Head of Kinsale, bordered by shore platforms. The coast is interrupted by Cork Harbour, a major ria formed by partial

marine submergence of a trellised network of valleys that were incised during Pleistocene low sea level phases, some following the east-west strike of the Carboniferous Limestone and Old Red Sandstone outcrops, others running transverse to it. Ballycotton Bay, Youghal Bay and Dungarvan Bay have been cut out of less resistant Carboniferous outcrops between headlands of hard Old Red Sandstone, as at Knockadoon Head. In Ballycotton Bay

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there are receding cliffs cut in glacial drift and Dungarvan Bayhas sedimentary infilling behind a barrier spit. To the east the coast is dominated by grassy bluffs with basal rocky outcrops, passing to higher cliffs on bold ­headlands. The Old Red Sandstone and Carboniferous Limestone give place to Lower Palaeozoic, mainly Ordovician rock formations. There are irregular modern shore platforms and dissected remnants of emerged shore platforms dating from Pleistocene interglacial high sea level phases. At Ardmore an erratic boulder on the shore platform is believed to have been ice-borne, released from an iceberg that became stranded here. It is difficult to explain this, as sea level was probably lower when the climate was cold enough for icebergs, but one suggestion is that they were emplaced here by a tsunami. There is a large dune-capped spit at the head of Tramore Bay, backed by a shallow embayment with broad mudflats exposed at low tide (Ruz 1987). Waterford Harbour is a ria at the converging mouths of the Suir, Nore and Barrow Rivers. On its eastern side is the limestone spur of Hook Head. To the east at Fethard are slumping cliffs in glacial and periglacial drift and Bannow Bay is a shallow tidal estuary at the mouth of Corock River. The Ordovician rocks end against Cambrian grits and shales (mainly Bray Group), interrupted by a synclinal corridor of Carbon­ ferous Limestone that runs NE froim Ballyteighe Bay to Wexford Harbour. A broad barrier spit impounds Bally­ teigh Lough, and to the east of Kimore Point a series of spits encloses Tacumshin Lake and Lady’s Island Lake, some with sand dunes and all showing signs of continuing modification by wave action (Ruz 1987). The extensive beaches and dunes have received sediment from outwash deposits at the limit of the Last Glaciation, which intersects the coast at the granite headland of Carnsore Point. Granite boulders are strewn along the shore from Kilmore Point east to Carnsore Point. The south-east coast of Ireland is generally low and hilly, with vegetated bluffs less than 30 m high fringed by beaches of sand and shingle derived from glacial drift. Wexford Harbour is a large open lagoon sheltered by spits of sand and gravel, Rosslare Spit, to the south and the wider Raven Spit to the north. The River Staney flows into the western side, and there was extensive reclamation in the nineteenth century on the northern and southern parts of this lagoon, behind the two spits. The spits have shown landward movement and occasional breaching in response to storm wave action. In southeast County Wexford there are extensive cliffs cut in glacial drift. Near Blackwater they intersect the Screen Hills kame-kettle moraine and the coast has

retreated considerably (Stephens 1970), but abandoned cliffs occur as vegetated bluffs in these glacial deposits where there has temporary accumulation of substantial foredunes. The coast is generally low-lying, but there are occasional bolder rocky headlands, such as Wicklow Head, cut  across ridges that follow the E-W regional strike of Cambrian, Ordovician, Silurian and Carboniferous formations. North of Wicklow Head a coastal lowland is fringed by a long shingle beach, The Murrough, ex­­tending to Greystones. At Bray Head cliff recession in Cambrian rocks has necessitated the realignment of a coastal railway. The seaside resort of Bray has a shingle beach which has been artificially renourished (>Fig. 7.26.13). The coastline of Dublin Bay has been much modified by the construction of sea walls, breakwaters and harbour structures, as at Dun Laoghaire on the southern shore. Long breakwaters border the mouth of the River Liffey, where Dublin Harbour has extensive docks. On the northern side Bull Wall runs out from Clontarf and forms an artificial southern boundary to a barrier island, North Bull Island. This stands in front of a tidal lagoon than opens to the north by Sutton Strand. The coast curves out to Howth Head on Cambrian rocks, a former island attached to the mainland by a tombolo. At Sutton a postglacial raised shingle beach 4 m above mean sea level rests upon a kitchen midden about 5,250 years old. Howth has wide sandy beaches, extensive sand shoals and harbour breakwaters, and Ireland’s Eye is a rocky island offshore, rising 69 m above sea level. To the north dune-capped barriers of sand and gravel enclose former embayments as estuarine lagoons, which have changed in configuration partly as the result of nineteenth century railway construction (Mulrennan 1993). Maynetown Lagoon is bordered by a barrier spit behind Sutton Strand, and opens to the sea between paired spits at Portmarnock Point and Cush Point, with extensive sand shoals offshore. At Malanide the railway embankment impounds a freshwater lake to the west, leaving the remains of an estuarine inlet with tidal channels to the east. The reef-bearing Carboniferous Limestone of the Dublin district gives place northward to the Ordovician sedimentary rocks and volcanic formations of Portrane and Lambay Island, and strongly-folded strata are well exposed in the cliff behind Loughshinny Harbour (>Fig. 7.26.14). At Skerries hard Silurian rocks outcrop on Red Island, attached to the mainland by a sandy tombolo, and the shingle beach west of the harbour has buildings armoured against storm wave damage. Balbriggan stands

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⊡⊡ Fig. 7.26.13 The shingle beach at Bray, on the east coast of Ireland, has been artificially renourished in front of the promenade. (Courtesy Geostudies.)

⊡⊡ Fig. 7.26.14 Cliff cut in strongly folded Ordovician rocks at Loughshinny, County Dublin. (Courtesy Geostudies.)

on an anticlinal outcrop of Ordovician slate, and the Boyne River estuary flows through a corridor of Carboniferous Limestone. Ordovician rocks return in the Clogher Head peninsula, where the shore rocks show dissection of sharp anticlines. The wide Dundalk Bay which has a broad intertidal shore with numerous sand bars and shoals along the shore, and beaches backed by dunes. Shelly beaches derived from the intertidal sandflats line the shore at Annagassan, and

wide sand bars and shoals are exposed at low tide off Blackrock. Sheltered embayments to the north are occupied by salt marsh, and Dundalk Harbour occupies a deepened estuary. To the north are cliffs and bluffs cut in glacial drift, which are steep and high at Rathcor, where a moraine is truncated and at Templetown Point, which has a capping of glacial outwash sand and gravel. They extend round into Carlingford Lough on the border with Northern Ireland.

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References Davies GLH, Stephens N (1978) The geomorphology of the British Isles: Ireland. Methuen, London Duffy MJ, Devoy RJN (1999) Contemporary process controls on the evolution of sedimentary coasts under low to high energy regimes: Western Ireland. Geologie en Mijnbouw 77:333–349 Guilcher A (1966) Les grandes falaises et megafalaises des cotes sudouest et ouest de l’Irlande. Ann Géorgr 75:26–38

Guilcher A, King CAM (1961) Spits, tombolos and tidal marshes in Conne­ mara and West Kerry, Ireland. Proc Royal Irish Acad, 61B:283–338 Mulrennan M (1993) Changes since the nineteenth century to estuarybarrier complex of north County Dublin. Ir Geogr 26:1–13 Ruz MH (1987) Recent evolution of the southeast barrier coast of Ireland. J Coastal Res 5:523–540 Stephens N (1970) The coastline of Ireland. In: Stephens N, Glasscock RE (eds) 1970 Irish geographical studies in honour of E. Estyn Evans. Queen’s University Belfast, UK, pp 125–145

7.27 Channel Islands

Eric Bird

1.  Introduction The Channel Islands (British dependencies, and otherwise known as Les Iles Normandes) lie about 130 km south of Portland Bill (Dorset) and to the west and northwest of the Cotentin Peninsula (Normandy, France). They include Jersey about 105 sq. km), Guernsey (62 sq. km), Herm (3 sq. km), Sark (5.5 sq. km), Alderney (8 sq. km) and a number of smaller rocky islets, such as Les Minquiers and the Ecréhou Rocks. They are shown on local topographic maps and British Geological Survey sheets 1 and 2. They consist mainly of Palaeozoic metamorphic and igneous rocks and are generally dominated by plateaux bordered by steep and cliffy coasts and wide rocky shores exposed at low tide. Sand and shingle beaches occupy bays, but erosion has been a problem, and since the eighteenth century there has been extensive construction of sea walls, mainly of dressed granite (Bird 2002). These have held the coastline, but resulted in further beach depletion by wave reflection scour, so that most beaches are now submerged at high tide (Gunton 1997). There has been some artificial beach replenishment, but generally the focus of interest for local engineers has been on maintaining and elaborating sea walls rather than restoring beaches. The climate is mild and maritime, with prevailing westerly winds and frequent winter storms. Mean annual rainfall is between 750 and 900 mm. Wave action is stronger on southern, western and northern coasts than on eastern coasts, but is impeded by the large tide range: St. Helier in Jersey has a mean spring tide of 9.8 m, and in places the mainly rocky intertidal zone is 2–3 km wide. The steep coasts show slope-over-wall profiles, the slope mantled by periglacial Head deposits above a rocky cliff cut by wave action during the period since the sea rose to its present level about 6,000 years ago. The proportion of cliff to upper slope increases with exposure, with some vertical cliffs on western coasts and Head-mantled slopes descending almost to high tide level on sheltered parts of the east coast.

2.  Jersey Just under 20 km west of the Normandy coast, the island of Jersey measures about 16 km W–E and 8 km N–S.

It consists of a broad plateau (> Fig. 7.27.1) that slopes southward from about 120 m near the north coast, and is incised by several sub-parallel valleys with rivers draining mainly to St. Aubin’s Bay on the south coast. St. Helier, the capital, stands at the eastern end of St. Aubin’s Bay, with a harbour protected by large stone breakwaters and the Elizabeth Marina built in 2000. At low tide a kilometrewide sandy shore is exposed, and an ebb road runs out along a causeway to the rocky island on which the sixteenth century Elizabeth Castle stands. To the east is the high promontory of La Collette, where ash waste from the power station has been used to reclaim coastal land. The coast to the southeast has sea walls overlooking a shingle beach and wide expanses of sand and rock outcrops at low tide. Accretion beside a slipway indicates predominant westward drifting. There are low rocky headlands at Le Croc, Le Nez and Le Hocq, and St. Clements Bay has a narrow beach behind a low tide rocky shore dissected by sand-floored channels cut out along major joints. At Plat Rocque Point a stone break water shelters a high tide harbour, and at low tide rocky reefs and sandy channels extend out more than 2 km to the southeast, with the Seymour Tower as a navigation mark. At La Rocque Point the coast turns northward into the Royal Bay of Grouville. Sea walls line a shore that becomes increasingly sandy, and give place to grassy dunes north of Le Hurel. The low-lying hinterland, Les Marais, was formerly a marshy lagoon. To the north the coast curves out to Mont Orgueil above Gorey, where a breakwater shelters another high tide harbour. The upland of Faldouet then comes to the coast as a series of rocky headlands between bays and coves that have grey shingle upper beaches and wide sandy shores at low tide. The shingle has been derived from weathering and erosion of the Lower Palaeozoic Ronez Conglomerate which outcrops in the north-eastern corner of the island. The bay at Anne Port has a shingle beach (> Fig. 7.27.2) which became depleted by reflection scour in front of the sea wall, and in 1999 was nourished with well-rounded grey and pink pebbles imported from a quarry near Cork in southern Ireland, chosen because of their similarity to the natural beach of pebbles from the Lower Palaeozoic Ronez Conglomerate. The shingle forms an upper beach which descends to a sharp border with

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_7.27, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 7.27.1 The plateau of northern Jersey, incised by valleys running southward. (Courtesy Geostudies.)

⊡⊡ Fig. 7.27.2 Anne Port, showing renourished shingle beach. (Courtesy Geostudies.)

low tide sand. It has been necessary to protect an eroding cliff on the southern side (Le Saute Geffray) with a gabi­ ­on  mantle to prevent a house and the coast road being undermined. Havre de Fer occupies a similar bay, with a jetty out to the Archirondel Tower on a rocky island, then shingle beaches descend to sandy and rocky shores in St. Catherine’s Bay to the north. The long St. Catherine’s Breakwater runs out ESE from runs out from Le Fort du Verclut and north of Verclut Point there are vegetated coastal slopes without

basal cliffing, extending behind Fliquet Bay to La Coupe Point, the northeast corner of Jersey. There are sandstones and slates associated with the Rozel Conglomerate. The rocky islets, Les Ecréhou, are 8 km out to the northeast and the Normandy coast 19 km east. The north coast of Jersey is steep and high, with slopes mantled by periglacial drift (Head deposits) descending to rocky cliffs and sloping foreshores that are much narrower at low tide than on the southern and eastern coasts. Shingle derived from the Ronez Conglomerate forms cove beaches,

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⊡⊡ Fig. 7.27.3 Shore platform on Rozel Conglomerate, Rozel Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.27.4 Steep coast of Bouley Bay, with basal ledges which are remnants of emerged Pleistocene shore platforms. (Courtesy Geostudies.)

with sand in Rozel Bay sheltered by a stone jetty. A sandy slope is exposed at low tide, wet with groundwater seepage, and there are brick enclosures (fish traps) on shore reefs of conglomerate which are partly wrack-encrusted and partly bare (> Fig. 7.27.3). To the east steep slopes ands basal cliffs run out to Nez du Guet and White Rock, painted as a sea mark. The steep embayed coast continues past L’Etacquerel, on which is a round tower and a ruined stone fort, to La Tete des Hougues, a craggy promontory where the Rozel

Conglomerates end. In Bouley Bay the rocky shore has a prominent black horizon just above normal high tide, overlain by orange lichens, then grassy, heathy, bracken or scrub vegetation on steep coastal slopes on the Bouley Rhyolites, harder parts of which persist as rocky tors. The low cliffs are fronted by ledges with are remnants of a Pleistocene shore platform (> Fig. 7.27.4), better formed than any existing intertidal shore platform, and clefts in the heathy slope descend to coves with very little beach material.

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An enclosed stairway descends to the cove at Porteret, which has a stone jetty and beaches of grey shingle above low-tide sand. The steep coastal slope, mantled with periglacial drift, is well preserved northward to Vicard Point, a headland of volcanic tuff and mudstone. And on past the little cove of Le Petit Port, which has shingle (derived from Rozel Conglomerate outcrops in slope tors) to La Colombière. Offshore the sea washes round the ugly rocks of Les Sambues. The cliffs becoming bolder above rocky Les Ruaux, where there are caves cut out along faults in a complex of granite and diorite, and a shelly raised beach below Belle Hougue Point. Gifford Bay (> Fig. 7.27.5) and Bonne Nuit Bay have cobble beaches, and a stone jetty protects a high tide harbour below very steep coastal slopes at the western end of Bonne Nuit Bay. From Fremont Point steep slopes, basal cliffs and rocky shores extend west along St. John’s Bay to Ronez Point, where there is a large coastal quarry in granite and diorite (> Fig. 7.27.6). Quarry waste has spilled on to the shore and drifted eastward to form gravelly beaches as far as Mourier Bay. The steep coast continues southwestward, with minor headlands and a series of bays and valleymouth inlets, as at Les Mourier where the river descends a waterfall. Sand is exposed at low tide in the bay at Les Reuses and Crabbé and Creux de Lasse are steep-sided inlets trenching the coastal plateau. A larger valley descends to La Greve de Lecq, where a stone jetty protects a little harbour, its foundation running out as a rocky reef. The jetty was truncated by a storm early in the twentieth

century. The beach is of pink sand, derived from the coarse-grained pink granite that outcrops in and around the bay (the granite of the north–west igneous complex). A stream pours out of a pipe in the wall, and has washed out a channel across the sandy beach in front of a roadside sea wall, where a dune and grassy terrace have been consumed by erosion in recent decades. The steep and cliffy coast now runs northwest past many small coves to Pièmont Point, and the bay to the west is La Greve au Lanchon, with a sandy shore exposed at low tide. Similar steep slopes continue to Gros Nez Point at the northwest corner of Jersey. The west-facing coast south from Gros Nez Point is steep with bold granite cliffs facing waves from the Atlantic Ocean. At Le Roudil de Puleq the cliffs of pink granite pass into bluffs of gneiss behind a small cove (> Fig. 7.27.7), the granite boulders on the shore giving place to grey cobbles below the gneiss. A rocky foreshore cut in Jersey Shales widens off the Grande Etaquerel promontory, which marks the beginning of St. Ouen’s Bay, where the steep coast passes inland behind a coastal plain. The sandy beach in St. Ouen’s Bay is over 5 km long and up to half a kilometre wide at low tide (Hydraulics Research 1991). It consists partly of quartzose sand from weathered granites and partly of calcareous biogenic sand from the sea floor. It is backed by a sea wall up to 4.5 m high (> Fig. 7.27.8), with some grassy dunes and reedy lagoons on a coastal plain backed by a bold escarpment. The coastal plain is dominated by Holocene dune sand

⊡⊡ Fig. 7.27.5 Shingle beach in Giffard Bay. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.27.6 Coastal quarry at Ronez Point. (Courtesy Geostudies.)

⊡⊡ Fig. 7.27.7 Cliffs of granite and gneiss, Le Roudil de Puleq. (Courtesy Geostudies.)

­ nderlain by fine angular solifluction gravel, and then u older (Pleistocene) beach and dune sand, as seen in Simon’s Sand and Gravel Pit. The sandy beach, exposed only at low tide, has been depleted by scour due to wave reflection from the sea wall. The natural supply of sand from alongshore and offshore is slow, and the dredging of shelly sand from the sea floor to the south–west (off Corbière) in the 1980s may have diminished the supply

from offshore. The beach is now low and almost flat, with extensive seepage zones, and its depletion has exposed outcrops of the underlying Jersey Shale, as well as some local underlying peat and fossil trees: a submerged forest. In the middle of the beach, near the Ouzière Slipway, large angular blocks have been dumped to protect the base of the sea wall and small concrete groynes built to give local access to the beach.

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⊡⊡ Fig. 7.27.8 Beach at low tide, St. Ouen’s Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 7.27.9 La Corbière. (Courtesy Geostudies.)

To the south the high sea wall is backed by grassy dunes that extend up and over the escarpment at Les Quennevais. The sea wall has been undermined, and in places is tilted and cracked, notably in the sector which was built during the German occupation in the Second World War (1942) to deter Allied invasion, rather than to halt coastline recession: it was built up to 50 m seaward of a dune cliff, and the intervening area was filled with sand

and gravel taken from the beach. Although there are some sectors of shingle along the base of the wall the response has been to protect it with further concrete ledges rather than by beach renourishment. The southern part of the beach is dominated by sand of granitic origin, the Jersey Shales giving place to extensive reefs of coarse-grained Corbière granite, which runs inland and is exposed in the quarried spur at La Carrière. The

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beach comes to an end against a high promontory at La Pulente, and the southwestern upland runs out to the long narrow peninsula of La Corbière, where the lighthouse stands on an outlying rocky reef at the southwest corner of Jersey (> Fig. 7.27.9). The steep coast turns eastward from La Corbière in a succession of small bays and headlands to St. Brelade’s Bay, where a wide sandy beach is exposed at low tide in front of a sea wall. A high promontory then runs out to Point La Fret and Noirmont Point, where the steep coast runs north along the western side of St. Aubin’s Bay. The coastal slopes descend to low cliffs on this more sheltered sector, then pass inland behind St. Aubin, which has breakwaters  defining a high tide harbour. The long gently curving coast of St. Aubin’s Bay is backed by a narrow coastal plain, across which several rivers flow to the sea. The coastline is walled, and a wide sandy shore is submerged at high tide.

3.  Guernsey Guernsey is dominated by Pre-Cambrian igneous and sedimentary rocks. Much of the island is a plateau about 90 m above sea level, bordered by steep coasts, but the northern part is low-lying with small outcrops of hard rock known as hougues and flat areas that were formerly marshy lagoons. Sand and shingle beaches occupy bays between low rocky promontories along the northwest coast, where wide intertidal rocky shores are valued as ⊡⊡ Fig. 7.27.10 St. Peter Port, Guernsey. (Courtesy Geostudies.)

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habitats for fish, shellfish and birds. At intervals around the coast there are emerged Pleistocene beaches 30 m, 18 m and 8 m above mean sea level, with associated loess and Head deposits. The east coast of Guernsey is relatively sheltered from the prevailing westerly winds and strong wave action from the Atlantic Ocean, and St. Peter Port (> Fig. 7.27.10) has developed in stages of breakwater construction which have attached the rocky island on which Castle Cornet stands. Mean spring tide range at St. Peter Port is about 9 m, and there is an intertidal rocky area up to 600 m wide on the coast to the north. The shore and coast consist of layered gabbro St. Peter Port stands on a slope rising to the plateau, but the lower country to the north is reclaimed marshland, commemorated by the Chateau des Marais. The coast north of St. Peter Port is lined with high sea walls to Salerie, and on round Belle Greve Bay. On the southern side of this bay there is no beach in front of the sea wall, but farther north a beach of grey shingle faces south–east. The wide intertidal rocky area is layered gabbro, dissected by channels, some of which have a sandy floor. The harbour at St. Sampson occupies the eastern end of a former strait that ran across to L’Islet on the north coast, separating the northern part of Guernsey as an island (Closs du Valle). The name Bridge persists at the inner end of the harbour, where the ancient bridge became a solid quay, and the strait has disappeared as the result of reclamation. There has also been land reclamation along the coast on either side of St. Sampson, and northward to Bordeaux Harbour.

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The gabbro that dominates the shore between St. Peter Port and Bordeaux Harbour gives place to diorite, which extends round the north coast. There is a diorite quarry near to the coast at Cockagne, and former granite quarries are now flooded at Millette Bay and the Beaucette Marina. Fort Doyle stands on a low diorite promontory, as does Fort Le Marchant to the west of Fontenelle Bay. L’Ancresse Bay has a wide sandy shore exposed at low tide, backed by segments of high tide shingle beach along a sea wall, part of which was constructed by the German occupation forces in 1942. The bay is backed by dunes which formed when sand was blown from the shore, but sea wall construction has been followed by disruption of the link between the beach and the dunes and lowering of the beach profile. The Mont Cuet peninsula is low-lying and fringed by diorite and granite shores. These rocks are overlain by a thin mantle of sandy drift (Head deposits formed under Pleistocene periglacial conditions) and then sand dunes. Sandy beaches, backed by grassy dunes, are extensive in Grand Havre, interspersed with outcrops of intricately jointed diorite and granite, from which boulders, cobbles and pebbles have been derived. In places the dune fringe was eroded, and is now protected by dumped granite boulders, as in Port Soif (> Fig. 7.27.11), but in others (as at L’Islet and in Ladies Bay) the grassy dunes extend uncliffed to the high tide mark. Portinfer has dunes behind a shingle beach. There are low cliffs cut in Head deposits on several of the west coast promontories, notably at Grande Rocque which is on granite. Saline Bay merges into Cobo Bay,

where a broad sand-floored inlet on the rocky shore allows larger waves to break against the sea wall at high tide. The sandy beach is wider, higher and drier on either side, in the lee of nearshore rocky reefs. Fort Hommet stands on a longer granite promontory, and to the south Vazon Bay has a wide sandy shore with patches of granite crossed by many dolerite dykes. The beach is lower behind a gap through the shore rock outcrops, and groynes inserted in the hope of trapping sand or shingle to protect the sea wall proved futile, and are now derelict. The problems of maintaining sea walls and fronting beaches here are similar to those in St. Ouen’s Bay in Jersey. Fort Richmond stands on a granite promontory with low eroding cliffs cut in Head and weathered rock, and Perelle Bay has a broad sand-floored channel between extensive granite shores where boats can be drawn up. Part of the sea wall is protected by an apron of shingle. At Saumare Bay a gravelly beach is backed by a ridge of well-rounded grey shingle built up by storm waves, which occasionally wash cobbles and pebbles over the crest. The ridge is kept in position by bulldozing overwashed material back to the crest. An artificial path runs across the broad rocky isthmus that emerges at low tide between Lihou Island and the mainland. There are paired curving shingle beaches on the inner shore of Lihou Island, and on the mainland coast. Lithou Island is also granitic, with veins of dolerite trending ENE and a capping of Head deposits. Rocquaine Bay has extensive areas of intertidal coarse-grained granite, much valued for shellfish and seabird habitats. There are segments ⊡⊡ Fig. 7.27.11 West coast of Guernsey beach, Port Soif. (Courtesy Geostudies.)

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of sand and gravel beaches in front of a high sea wall, notably in Portelet Harbour at the southern end. The beach has been lowered by wave reflection in front of the sea wall, and when it is exposed at low tide there is extensive seepage of water draining beneath the wall from the hinterland. At Portelet Harbour the steep coast begins, with a slopeover-wall profile in which the slope is mantled by periglacial drift deposits (Head) which outcrop above granite exposed in the basal cliff. This extends round the southwestern headland, past La Pointe de Pleinment, and on along the south coast of Guernsey. The granite gives places to metamorphosed granite gneiss which dominates the south coast and extends up the east coast to St. Peter Port. There are many small bays and inlets, and the rocky shore is diversified by clefts and caves. Beyond Icart Point the slopeover-wall coast runs past Saints Bay, a valley-mouth cove, and behind Moulin Huet Bay, where a broad sandy beach is exposed at low tide and there are small high tide shingle beaches in coves. On the eastern side, above Petit Port, the coastal slope is interrupted by tors and ­buttresses of hard quartzite, which persisted when the ­surrounding granite gneiss was shattered by periglacial processes (> Fig. 7.27.12). Jerbourg Point is a jumble of rocky tors and ­buttresses on a Head-mantled slope undercut by rugged cliffs, descending to the Pea Stacks. Ribs of harder rock run out across the shore, and there is a shingle beach at high tide in Vaux Beres. St. Martins Point is a long protrusion of steeply-dipping rock, culminating in an island linked by a bridge, marking the southeast corner of Guernsey.

⊡⊡ Fig. 7.27.12 Coastal tors, Moulin Huet Bay. (Courtesy Geostudies.)

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The coastal slope of southern Guernsey carries a heath or scrub vegetation, but beyond St. Martins Point, where it swings northward and becomes more sheltered, this passes into woodland. Just north of St. Martins Point is a cauldron in the coastal slope formed where a cave roof has collapsed. Marble Bay is a sheltered cove, beyond which Bec du Nez protrudes, then Fermain Bay has a wide beach of grey shingle at high tide. There were numerous small landslides in Head deposits and weathered gneiss on the coast at and north of Fermain Bay in the wet winter of 2000–2001. Just north of Fermain Point is a shore platform cut in diorite. The steep coast descends to a rocky intertidal area which is much narrower than that to the north of St. Peter Port. There are numerous clefts, caves and small coves, many cut out along dolerite dykes that outcrop in the cliffs. Soldiers Bay is a cove with a high tide beach of grey shingle and some sand, and the steep wooded coast, with slopes descending to basal cliffs of granite gneiss, continues round to Havelet Bay, where South Beach is a clean sandy beach above intertidal muddy sand flanked by South Quay at St. Peter Port.

4.  Herm Herm, 5 km east of Guernsey, is about 2.4 km from N–S and 0.8 km E–W. The southern part of the island is a plateau, rising southward from 40 to 60 m, and bordered

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by steep slopes, mantled by periglacial Head deposits, below which are rocky shores. The island has an area of about 200 ha, but is bordered by another 400 ha of intertidal rocky reefs and islets, with some high tide sandy beaches, linking sand bars and an extensive sandy sea floor. The northern part of the island is low-lying, bordered by grassy sand dune topography. From the landing pier at Rosière Steps on the west coast steep vegetated coastal slopes extend southward, mantled with periglacial Head deposits and breached by coves where caves have collapsed, as at Bishops Cove (> Fig. 7.27.13). Offshore is rounded Jethou Island, also granodiorite, where the vegetated Head-mantled coastal slopes descend almost to high tide level on the sheltered eastern side. From Point Sauzebourge on the southwest corner of Herm the south coast has higher cliffs behind rocky and boulder-strewn shores, the cliffs cut by deep chasms at Le Creux Pigeon and Barbara’s Leap. There has been local slumping and cliffing in the earthy Head deposits, and the granodiorite shores are intersected by numerous clefts. On the east coast there is only minor cliffing north from Moulinet, the Head-mantled vegetated slopes descending to low cliffs in earthy gravel and rocky shores, with the island of Caquorobert linked to Herm by an intertidal bouldery isthmus. The coastal slope is cut by little valleys with lush vegetation. Belvoir Bay is sandy, and the steep coast slopes down to a rocky granitic shore. The steep granitic coast ends at Les Ancelées. To the north is east-facing Shell Beach, which consists of shell grit and granitic granules with a scatter of whole shells,

mainly Patella, from the shore rocks and numerous small shells including species evidently from the Caribbean: it has been suggested that they came by way of the Gulf Stream. It is not clear why shells are more abundant here than on other Channel Islands or Cotentin beaches. They have in any case been depleted by collectors. La Pointe du Gentilhomme, the north–east corner of Herm, has numerous granodiorite islets and reefs offshore, some with high tide sand beaches, others linked by sand bars, and the shallow sea floor is sandy. Mouisonnière, a north-facing sandy beach (> Fig. 7.27.14), is backed by cliffed dunes, and on the west coast, south from Oyster Point, the Bear’s Beach and Fisherman’s Beach are also sandy. Herm is sheltered by Guernsey from westerly gales and storm waves from the Atlantic, and these west-facing sandy beaches are generally stable, backed by gently sloping grassy dunes. To the south the foreshore becomes rocky and littered with boulders as the steep granitic coast begins, and an intertidal causeway runs out to a bouldery islet. Low cliffs back a beach of cobbles and boulders with some sand south to Rosière Steps.

5.  Sark The rocky island of Sark, 11 km east of Guernsey, consists of Great Sark, to the north, linked by a narrow 90 m high knife-edged rocky isthmus, La Coupée, to Little Sark, to the south, while Brecqhou is another small rocky island to the west. Foliated granodiorites dominate, with numerous ⊡⊡ Fig. 7.27.13 Collapsed cave on the southwest coast of Herm. (Courtesy Geostudies.)

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⊡⊡ Fig. 7.27.14 Cliffed dune at Mouisonnière, north coast of Herm. (Courtesy Geostudies.)

minor intrusions. Steep rocky embayed coasts rise to an undulating plateau up to 114 m above sea level, with cliffs dissected by many coves and inlets, caves and natural arches related to cleavage planes and minor intrusions. The shores are rocky with boulders and some shingle, but sand is exposed at low tide in a few of the coves, notably Derrible, Dixcart and Grande Grève.

6.  Alderney Lying 16 km west of Cap de la Hague on the Cotentin Peninsula in France, Alderney (5 km  ×  3 km) has steep coasts bordering a plateau about 60 m above sea level. The western third consists of foliated granodiorite and gneiss, the central section diorite, and the eastern fringe reddish brown Alderney Sandstone, possibly Lower Palaeozoic, with a basal conglomerate. The conglomerate is up to 2 m thick beneath a south coast outlier, where it yields pebbles to adjacent beaches. There are bold cliffs up to 83 m high on the exposed W and SW coasts, but otherwise low cliffs

back extensive intertidal rocky shores with many small coves and inlets, often related to the numerous felsite dykes. There are emerged Pleistocene beaches and terraces 25–30 m, 18 m, and up to 8 m above mean sea level, the 8 m beach overlying periglacial Head in Longis Bay. Sandy beaches backed by grassy dunes occupy the larger bays, as at Corblets Bay on the north coast, Braye Bay where a harbour is sheltered by a long breakwater, and Longis Bay on the south coast. Although the mean spring tide range is smaller (about 6 m) than in Jersey and Guernsey there are strong tidal currents in the Raz Blanchart, the strait between Alderney and Cape de la Hague.

References Bird ECF (2002) The sea walls of Jersey. Jersey Soc Bull 7/6:6–10 Gunton A (1997) Upper foreshore evolution and sea wall stability, Jersey, Channel Islands. J Coastal Res 13:813–821 Hydraulics Research Station (1991) Jersey coastal management study. Report EX2490. Wallingford, UK

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8.0 Europe – Editorial Introduction Europe is a peninsula projecting NW from the continent of Asia, lying to the west of the Ural Mountains, which run from north to south through Russia. There are four major geological divisions: the NW Highlands extending through >  Norway and > Sweden to the north and west of the >  British Isles; the Great Northern European Lowland, extending from > Russia westward to >  Germany, >  France and > Britain; the various Pre-Cambrian and Palaeozoic massifs, of which the Iberian Massif, Brittany, >  Corsica and > Sardinia extend to the coast; and the Alpine Mountains, including the Alps, Pyrenees, Apennines, Carpathian, and Sierra Nevada in southern Spain (Ager 1980). Each of these reaches the coast, and influences coastal morphology (Kelletat and Scheffers 2005). The coast is treated in sequence from > Norway in the north, round the Baltic and down the Atlantic coast, then through the northern Mediterranean into the western Black Sea, ending with the > Sea of Azov. The NW Mountains were originally formed by the Caledonian orogeny, folding and faulting Lower Palaeozoic formations into structures with a dominant NE-SW trend. The west coast of Norway follows this trend, although the fiords are glacial troughs excavated along valleys that run transverse to it, and its influence is seen in the coastal outlines of western Scotland and Wales. To the east of the Norwegian mountains is the broad Fenno-Scandian Shield, consisting of Pre-Cambrian igneous and metamorphic rocks prominent on the east coast of Sweden and around Finland. Along its southern margin are outliers of Lower Palaeozoic rock that are horizontal or slightly folded, in contrast with their strongly folded equivalents in the NW Mountains. To the south, in > Poland, > Germany, > Denmark and eastern > Britain, Mesozoic and Tertiary formations are gently folded, and covered by glacial drift deposits. Cliffs on the southern coast of the Baltic Sea and on either side of the North Sea expose segments of the Mesozoic and Tertiary rocks, or are cut into glacial drift deposits. South of the limits of Pleistocene glaciation (which extended from South Wales through the London Basin and across north Germany and Poland into Russia) the gently folded Mesozoic formations form scarps and vales, while Tertiary sediment occupy basins. These formations outcrop along the north and west coasts of France, interrupted by the Brittany massif.

Further south these formations become strongly folded in the Alpine zone. The Pyrenees extend to the coast of SW France and northern Spain, and to the Mediterranean coast across the Franco-Spanish border, where the component formations are exposed in cliffs. Similar folded structures are seen on the southern coast of Spain, but much of the coast of the Iberian Peninsula borders the crystalline massif and its marginal sedimentary basins. Alpine structures also intersect the Mediterranean coast of France east of Marseille, and influence then outlines of Italy and Sicily, and the east coast of the Adriatic Sea, where the long ridges and straits of the Dalmatian coast follow the Alpine trend. Similar structures extend to the coasts of Greece and Crete, European Turkey, Bulgaria and Rumania, and across the Crimean Peninsula. Sectors of the Great North European Lowland then extend to the coasts of the Ukraine and the Sea of Azov. Europe spans a latitudinal range of 45°, with climatic zones varying from the arctic north, where Svalbard has glaciers and periglacial processes are active in northern Scandinavia, to the Mediterranean, where warm conditions favour salt weathering, bioerosion, algal and vermetid encrustations, calcareous biogenic beach sediment and the formation of beach rock. In between is the temperate zone, with the effects of past glaciation and periglaciation evident as far south as Portugal. There are contrasts between the Atlantic coasts, exposed to ocean swell and storm waves, and the lower wave energy of the Med­ iterranean and Black Sea. Mean spring tide range is relatively small (generally < 2 m) along the Norwegian coast and negligible around the Baltic, but it increases southward from Denmark to the German, Dutch and Belgian coasts (5.4 m an Antwerp) and remains large (5–8 m) along the north coast of France. Spring tides exceed 10 m in the Baie de St Michel, and diminish westward to 5.7 m at Le Conquet.

References Ager DV (1980) The geology of Europe. McGraw-Hill, London Kelletat D, Scheffers A (2005) Europe, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 452–462

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.1 Norway

Tormod Klemsdal

1.  Introduction Only a few works on the geomorphology of the coast of Norway (>Fig. 8.1.1) have been published, and most of them have dealt with fiords and the strandflat. Reusch (1894) introduced the term strandflat and discussed its form and origin, and Nansen (1922) made a substantial contribution to the subject. The Hardangerfiord was treated by Holtedahl (1975), who also dealt with the strandflat and the bankflat of the coast of Møre (Holtedahl 1955, 1960). The strandflat was considered in a geomorphological study of the Lofoten-Vesterålen area by Møller and Sollid (1973). Most descriptions of the Norwegian coast are based on general geomorphological features and landscape types. In addition, shore zone and coastal features have been mentioned in different works dealing with Quaternary geology. The steep western margin and the more gentle eastern slope of the Scandinavian land block have strongly influenced the type of coast. Further, the zone of wave attack, the shore, is in Norway composed mainly of bedrock, which makes it necessary to deal with both coastal and shore zone morphology (Klemsdal 1979). The Nor­wegian coast is classified in terms of a combination of the morphology of the shore and the coastal land that borders it (>Fig. 8.1.1) (Klemsdal 1982). Along the Norwegian coast much of the shore consists of gently and steeply sloping ice-smoothed rocky outcrops and stony beaches. In some places shore platforms have been cut by abrasion, while sandy beaches and clayey shores are found locally. Frost shattering has produced spiky forms on rocky shores. Gently sloping ice-smoothed rocky shores have a great variety of forms, from the smallest roches moutonnées (only a few square metres in size and a metre high) to the largest (several tens of metres in both length and height), plus a combination of more or less gently sloping icesmoothed rocky slopes with some plastic scouring forms engraved on them (>Fig.  8.1.2). There are also steep or nearly vertical shores, ice-smoothed or glacially scoured and plucked, plunging into the sea. Abrasion shore forms are found along small stretches where marine abrasion, aided by frost action, has been

able to transform the previously glaciated landforms. Abrasion has produced such forms as notches, stacks, geos, and shore platforms in front of cliffs (>Fig. 8.1.3). Only on some parts of the coast Quaternary sediment, mainly till (glacial drift), have been reworked by waves. Winnowing of the finer particles has left the larger ones behind, the finer material accumulating in sheltered areas as beaches. These include boulder beaches, with residual boulders generally derived either in till in front of a moraine cliff, or from ground moraine dipping gently into the sea (>Fig. 8.1.4). There are also stony beaches, either erosional, with large cobbles rounded by wave action, or depositional, with smaller, well-rounded pebbles, often forming a shingle beach ridge. A combination of boulder and cobble beaches is the most common beach type along the shores of Norway. Sandy beaches are found where sand, chiefly washed out of till or fluvioglacial deposits, has accumulated, or been moved by longshore drifting to more sheltered shores. The longest continuous sandy beach on the Norwegian coast is only 4 km, most sandy beaches being only pocket beaches 50–200 m long (>Figs. 8.1.5 and > 8.1.6). Where the shore deposits consist of silt and clay they form a gently slope extending above and below sea level. In Norway these are sometimes termed “clayey beaches”, but generally they are known as tidal mudflats, found in sheltered areas where a small stream has brought suspended fine-grained sediment material to the shore, where a plain of clay gently emerges from the sea, or in a tidal area as an accumulation of the finer material washed out of nearby Quaternary deposits. The mean tide range is less than 0.3 m south of Utsira (>Fig. 8.1.10), increasing to about 1 m on the west coast and 1.5 m farther north, with a maximum of 2 m at Vardø in the northeast. Intertidal mudflats and salt marshes are found only locally.

2.  Coastal Types A morphogenetic classification of the Norwegian coast may be based on a combination of shore features and backing landforms (>Fig. 8.1.1).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.1

Norway

⊡⊡ Fig. 8.1.1 The coastal g ­ eomorphology of Norway. 1. STRANDFLAT COAST Divided in sections due to the width of the supramarine zone/submarine zone (the supramarine zone includes the skerry zone) 4a a. Broad / very broad 1d b. Broad / broad c. Broad / narrow d. Narrow / medium broad 4a 4b e. Narrow / broad 1d

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T. Klemsdal 2008

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5 FLAT ABRASION COAST 6 MORAINE TOPOGRAPHY COAST 7 MORAINE CLIFF COAST 8 SANDY BEACH COAST

Norway

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⊡⊡ Fig. 8.1.2 The rocky shore of the fjärd coast of the Hvaler archipelago is cut in the Pre-Cambrian Østfold granite. The gently undulating topography was formed by pre-glacial weathering, and Quaternary glaciation has formed the ice-smoothed surface.

⊡⊡ Fig. 8.1.3 The cliff and shore platform on the west coast of Søndre Søster in the Hvaler archipelago. The bedrock is a Permian lava conglomerate with numerous fissures and joints, which have guided frost weathering and wave erosion. A gravelly beach is seen in front of the cliff.

The strandflat (>Fig. 8.1.7) is a gently sloping bedrock plain extending above and below present sea level, locally mantled with sedimentary deposits. It is limited both landward and seaward by steeper slopes. The shore of the strandflat consists mainly of gently or steeply sloping ­ice-smoothed rocky shores and stony beaches. Sandy beaches or intertidal mudflats are found only in places with an abundance of loose material. The coastal plain has

a ­rugged terrain with small relief and dips gently into the sea. It produces an irregular coastline with very many bays, coves, inlets, headlands, and promontories, which, together with islands, islets, and skerries are the elements of the skerry zone. The bottom topography is the continuation of the terrain of the skerry zone and the supramarine part. The gradient of the emerged part varies between 5 and 25 m per km, and in general the strandflat

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⊡⊡ Fig. 8.1.4 The fossil moraine cliff on Jæren, produced by wave attack, is fronted by a broad beach of boulders and gravel of varying size.

extends 40 m above and 40 m below sea level, though in some areas it extends 100 m above sea level. The strandflat is found from the northern parts of Jæren to the western parts of Finnmark with a narrow bench, more like an abrasion platform, extending to ­eastern Finnmark. It can be classified according to the width of the emerged and submerged parts (>Fig.  8.1.1). The maximum width of the strandflat is about 40 km in the area of Hitra on the Møre coast, but generally the strandflat is about 16 km wide. The emerged part is normally 5–10 km wide, but can reach 15 km. Locally it diminishes to a few hundred metres, but even there it is a very important landform, as a large part of the population lives on it. The origin of the strandflat began when the land was exposed to denudation through the Mesozoic and the early Tertiary. In a warm climate with dry and wet periods, denudation produced a land surface, the paleic surface (Gjessing 1967), with well-rounded, subdued landforms in the peripheral parts of the land block. In the Tertiary the Scandinavian land block was elevated and tilted, giving a steeper slope toward the west and northwest and bringing the paleic surface to varying heights above sea level. Along the coast it varies from sea level to an altitude of 500–700 m. Probably coincident with this uplift, the climate became temperate and more humid. Fluvial processes dissected the paleic surface, particularly in the west and northwest. Along the coast marine abrasion and denudation of the paleic surface, may have initiated a peneplain. Cirque glaciers, descending from higher coastal mountains on to this planed surface along the coast, were ­important in the widening and the dissection of the

strandflat. Ice-currents of the inland ice and valley glaciers  from inland, spreading out in the coastal areas (Holtedahl 1929), may also have taken part in the devel­ opment. In interglacial times and at the beginning of ­glacial periods, frost weathering and marine abrasion were very rapid, and together with mass movement and littoral  transport, resulted in the evolution of the modern strandflat. Fiords are arms of the sea stretching inland between distinct glaciated slopes, which either plunge steeply into the sea or descend to a more gently sloping valley bench. These slopes continue down to the floor of the fiord, ­giving the general U-profile in cross section. The longitudinal profile consists of thresholds and troughs. The fiords are cut into a paleic plateau-like or a mountainous landscape of varying heights above the fiord. There are variations the relief and length of fiords (>Fig.  8.1.8). The fiords are ­generally bordered by steeply sloping rocky shores and stony beaches. Where rivers enter a fiord, there are deltas with stony and sandy beaches and sometimes intertidal mudflats. Sognefjorden is the longest (200 km) and deepest (1,308 m) fiord in Norway. Its width is 1–8 km. and the surrounding plateaus and mountains rise from 500 m above sea level in the western parts to 1,500–1,700 m above sea level in the eastern parts. Examples of the smallest fiords are found in eastern Finnmark and in the southernmost part of the country. The Varangerfjord and the fiords of the middle parts of Finnmark, the Trondheimsfjord and the Oslofjord, are all very wide in relation to the height of the surrounding land, and thus lack some of the typical fiord characteristics.

Norway

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⊡⊡ Fig. 8.1.5 The coastal g ­ eomorphology of Jæren, southwest Norway. The numbers in the legend refer to the numbers in >Fig. 8.1.1.

Coastal landforms of Jæren Southwest Norway

Tunganee

Stavanger Tananger

Solasanden

Gansfjorden Fossil beach ridge Sandnea

Skarasanden Brusanden Revesanden Revtangen Orresandan Nærlandeanden

Stoney beach Rocky shore of the Strandflat (1)

Flggjo River

Rocky shore along the fjord (2)

Orre Lake

Rocky shore of the fjärd coast (3) Moraine topography coast (6), with a beach of boulders and stones or a clayey beach

Hä River

Moraine cliff coast (7) (fossil) with a beach of boulders and stones in front

Brusand

Sandy beach coast, (8) with sand dunes behind the beach and smaller dunes of silt and a thin layer of this finest alloan material (silt) further Inland

Ogna 10 km T.Klemsdal 2008

A pre-existing river valley system directed the ice movements, and glacial erosion produced large U-shaped valleys and overdeepened troughs, most prominent on the western and north-western parts of the land block. Invaded by the sea, these valleys and troughs became the fiords. The Caledonian structure, mainly NE-SW, secondary NW-SE and N-S, is responsible for the pattern of the valley system and the fiords. The imprint of the structure along the same directions is also present at the strandflat. The

form of the Oslofjord is due to less resistant down-faulted rocks, Permian faulting zones, and ice convergence. As distinct from fiords, fjärds are found where an undulating land surface with subdued valleys slopes gently into the sea, making an uneven coastline with numerous islands and islets with headlands and coves. The narrow inlets are not very deep, and are bordered by a terrain of low relief (>Fig.  8.1.2). As on the strandflat, the skerry zone of the fjärd coast has a shore zone with ice-smoothed

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⊡⊡ Fig. 8.1.6 The sandy beach of Revesanden. The waves have sorted the Lista Stage moraine and accumulated sand on the beaches of Jæren, south from Revtangen. Behind the sandy beach is a zone of sand dunes, the highest reaching 15 m above sea level.

⊡⊡ Fig. 8.1.7 The Strandflat at Myre, Vesterålen, North Norway, here covered by an extensive peat bog.

rocky shores and stony beaches, with sandy and clayey beaches in smaller bays. The fjärd coast of southern Oslofjord and the Skagerrak, the continuation of that of the west coast of Sweden, has a rather varied and complicated pattern of fjärds because of the Precambrian crystalline rock structure. This structure was rejuvenated in the Permian and – in part, probably – in Tertiary times. In the inner part of the Oslofjord, the headlands, bays, and islands are due to lithological variations and structural contrasts produced by Caledonian folding of Cambro-Silurian sedimentary rocks. On the

eastern, gently sloping part of the land block with subdued paleic landforms, the ice movement was free and undirected. Glacial erosion smoothed the surface. Along zones of weakness small depressions and fissure valleys were formed among knolls and hills, which later, when the sea entered, became the fjärd coast. On the Norwegian coast, abrasion has formed rugged cliffs that descend from the gently undulating paleic surface to the sea. The cliffed coast may descend straight into the sea, with abrasion acting directly on the base of the vertical wall; or there may be an abrasion shore platform

Norway

8.1

⊡⊡ Fig. 8.1.8 The Lysefiord fiord from Prekestolen (the Pulpit), 600 m above sea level.

⊡⊡ Fig. 8.1.9 The high cliff north of Kjøllefiord truncates a gently undulating paleic (Preglacial) land surface. In some places the vertical cliff plunges directly into the sea; in others it is fronted by a talus of loose material.

in front of the cliff. Frost weathering and marine abrasion in zones of weakness within the rock outcrops are responsible for the forms of the rugged cliff. The height of the cliff varies, and may attain 200 m (>Fig. 8.1.9). Smaller cliffs are found at various places along the coast, but the most famous localities of cliff abrasion coast are at Stad and the North Cape. From eastern Troms to eastern Finnmark, there are many localities with a cliffed coast. On the cliffed coast of Åna Sira, marine abrasion has

been of little consequence, and the cliff may be a result of Tertiary faulting, which also is a contributing factor at Stad and along the Finnmark coast. Waves cause marine abrasion, and have been recorded by lighthouse keepers only through subjective evaluation. Recently due to the need for better wave data in connection with petroleum installations and research on wave-­ generated electric energy progress has been made in registration of height and direction of waves. The present

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⊡⊡ Fig. 8.1.10 Factors involved in coastal development in Norway. The coastline displacement curves are in radiocarbon years bp.

Norway

data are sufficient to describe a yearly variation of the wave climate, the combined distribution of height, period, and direction of the waves through the seasons, along the coast. There are lower waves in the summer than from October through February and the data for March, April, and September are close to the mean value for the year. This variation in frequency of the significant wave height is demonstrated by the conditions at Utsira (>Fig. 8.1.10), which also shows the variation of significant waves along the coast. The area around Stad and Kråkenes has the largest waves and supply of wave energy, which decrease both northward and southward to nearly the same value at the Oslofjord and Vardø. Variations in wave climate are, however, less important to the effect of marine abrasion than are the properties of the rock, as indicated by the abrasion along the coast of Finnmark. There are sectors of flat abrasion coast, which consists of a bedrock plain concordant with the structure of slightly metamorphosed sedimentary rocks sloping gently into the sea. Marine abrasion of the rocks, accompanied by frost weathering, has produced a rocky abrasion shore with minor stony beaches. This type of coast is restricted to the north side of the Varangerfjord and as far east as Vardø, partially mapped by Tolgensbakk and Sollid (1980). Morainic coasts are found only on parts of Lista and Jæren (>Fig. 8.1.1). They result from an uneven ground moraine terrain of low relief, giving a coastline with headlands and bays. This type of coast has a stony beach with some boulders, little affected by the waves, in between well-rounded stones, with finer particles washed away. Cliffs have been cut into morainic coasts on parts of Lista and Jæren (>Figs.  8.1.4 and > 8.1.5). They form slopes 10–15 m high, with an inclination of 10–15° and locally steeper. The cliff has become a bluff covered with vegetation and elevated a few metres above the attack of the sea by isostatic rebound; it is only a few tens of metres away from the present shore. Boulders and cobbles derived from the till make a stony beach, the most important beach type on morainic coasts (>Fig. 8.1.4). Sandy beaches are found on Jæren at Skarasanden, Bore, Revesanden at Revtangen Orre, Nærland, Brusand, and on Lista at Kviljo. There are also long sandy beaches

8.1

on Vigra, an island on the Møre coast, and on Andøya. Many small sandy pocket beaches are found along indented rocky coasts. Some sandy beaches are backed by gravelly beach ridges, others by dune fields (>Fig. 8.1.6) (Klemsdal 1969). Beaches are shaped by incident waves, the direction of which depends largely on onshore winds, as modified by coastal configuration and nearshore topography. Much sand has been derived from glacial and fluvioglacial drift deposits, with some coming in from the sea floor, but rather little from present-day rivers. In addition to beaches along the present coastline there are emerged beaches resulting from postglacial isostatic–eustatic movements. The coastline displacement curve (Hafsten 1979) is shown in (>Fig. 8.1.10). The highest marine limit of Late Glacial time varies from 220 m above sea level in the Oslo area to only about 10 m at Lista and Jæren: where the Holocene marine limit is low wave attack has been concentrated on the coast at a nearly constant level. In the Oslofjord area the relative positions of land and sea in the beginning of postglacial time were not stable long enough to allow littoral forms to develop, but in Finnmark some excellent emerged beach ridges are found.

References Gjessing J (1967) Norway’s paleic surface. Nor Geogr Tidsskr 21:69–132 Hafsten U (1979) Late and Post-Weichselian shore level changes in South Norway. In: Oeke E, Schuttenhelm RTE, Wiggers AJ (eds) The quaternary history of the North Sea. pp 45–59 Holtedahl H (1955) On the Norwegian continental terrace, primarily outside Møre-Romsda1: its geomorphology and sediments. Bergen Universitet Årbok Naturvitenskap 14:1–209 Klemsdal T (1969) Eolian forms in parts of Norway. Nor Geogr Tidsskr 23:49–66 Klemsdal T (1979) Coastal, shore/beach and eolian geomorphology. proposal for terminology. Nor Geogr Tidssk 33:159–171 Klemsdal T (1982) Coastal classification and the coast of Norway. Nor Geogr Tidsskr 36:129–152 Møller JT, Sollid JL (1973) Geomorphological map of Lofoten-Vesterålen. Nor Geogr Tidsskr 27:195–205 Nansen E (1922) The strandflat and isostasy. Videnskapsselskapets Skrifter, I Mathematisk Naturvitenskapelig klasse 11:313 pp Tolgensbakk J, Sollid JL (1980) Vardø, geological and geomorphological map-(1:50,000). Geografisk Institutt, Universitetet i Oslo

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8.1.1 Svalbard and Jan Mayen

Tormod Klemsdal

1.  Introduction The Svalbard archipelago lies about 700 km north of Norway between 74°N and 81°N and between 10°E and 35°E (> Fig.  8.1.1.1). It consist of six large islands: Spitsbergen (39,043 sq. km), Nordaustlandet (14,210 sq. km), Edgeøya (5,030 sq. km), Barentsøya (1,330 sq. km), Prins Karls Forland (640 sq. km) and Kong Karls Land (326 sq. km) – and a great number of smaller islands. It also includes Bjørnøya (Bear Island) (178 sq. km), 440 km north of the mainland of Norway and 250 km south of Svalbard. The Svalbard Treaty (1925) gave custody of these islands to Norway. The geology of Svalbard and Bjørnøya (Winsnes 1988)  includes the Hecla Hoek Complex, found in the northern part of the Nordaustlandet, in the northeast and along the western coast of Spitsbergen, on Prins ⊡⊡ Fig. 8.1.1.1 The Svalbard archipelago with names of the larger islands and places mentioned in the text and a geological map.

Karls Forland and in the southern part of Bjørnøya (>Fig.  8.1.1.1). The Hecla Hoek Complex consists of mainly Pre-Cambrian metamorphic rocks with some Cambrian and Ordovician sedimentary formations. The  rocks of the Hecla Hoek Complex were metamorphosed during the Caledonian orogeny in late Silurian – early Devonian, and varies from highly metamorphosed gneisses and migmatites to marble and schists, quartzites  and limestone/marble. North of Kongsfjorden the Complex consists of gneisses and migmatites with Cale­ donian granites and basic intrusives, which are also found, together with igneous rocks like quartz porphyry lava, on the north eastern parts of Nordaustlandet. The younger sedimentary rocks vary. The Devonian rocks are shale and sandstone, while Carboniferous and Permian rocks consist of continental sandstones, beneath limestone and chert with seams of coal. Clay, shale and siltstone, and younger sandstone of Triassic age are followed by Jurassic shale and conglomerates. Cretaceous sandstone and shale underlie the sandstone, and there are Tertiary shales and seams of coal. The Devonian rocks are found in a rectangular area in the central, northern parts of Spitsbergen (>Fig. 8.1.1.1). The Carboniferous and Permian rocks occur in the southern parts of Nordaustlandet, in the north-eastern and central parts of Spitsbergen and in small areas in the western parts of Spitsbergen, from west of Ny-Ålesund in the north to the area of Hornsund in the south. The Tertiary rocks constitute the rocks of the central parts of southern Spitsber­gen, south from Isfjorden. Around these Mesozoic rocks are found in the central and southern parts of Svalbard. The islands of Barentsøya, Edgeøya, Svenskøya, Kongsøya and Hopen all have Triassic sedimentary rocks intruded by dolerite of Mesozoic age. The Caledonian Orogeny caused folding, thrusting and faulting of the bedrock. The opening of the North Atlantic Ocean in Tertiary times, with faulting and flexuring of the bedrock produced the present geological setting. Meteorological stations on Svalbard (> Fig.  8.1.1.1) are at Ny-Ålesund, Isfjord Radio, Longyearbyen, Svea­ gruva, Hopen and Bjørnøya. Mean annual precipitation varies between 300 and 500 mm along the coast, and the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.1.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Svalbard and Jan Mayen

average annual temperature is close to −6°C. The mean monthly temperature ranges from −15°C to 6°C, and is above freezing between the beginning of June and the end of September. As frost weathering is most active with temperature variations around 0°C to −5°C frost weathering along the coast is most active from April to June and from September to December. Wind-generated waves strongly affect the shore. The strongest winds are from SW, across open sea with a large fetch from the North Atlantic Ocean. Autumn is when the strongest winds produce the largest waves attacking the coast. Evidence from satellite images has enabled the Nor­ wegian Meteorological Institute to make maps of the average monthly extent of sea ice along the coast of Svalbard. Sea ice starts to form along the northern and northeastern coast of Svalbard in October and by December the archipelago is surrounded. The sea ice breaks up sufficiently for waves to attack the shore again in late May or early June on the southwestern coast of Spitsbergen. On the west coast of Spitsbergen sea ice is melted by the warm North Atlantic current before the rest of the coast is free. The west coast of Spitsbergen is ice free for an average of 5 months, while the coasts of Edgeøya and Barentsøya have a 3 month period free of ice from mid-July to the end of October. However, drifting sea ice can make it impossible for ships to sail around the archipelago during the summer, and when sea ice is driven onshore it erodes cliffs and rocky outcrops and moves beach material, including large boulders. Permafrost also affects the shore. During the summer months the upper part of the permafrost thaws, forming a layer of unfrozen ground that can be attacked by waves. When the temperature falls below zero and the ground freezes again the freezing of the water in the zone between the permafrost and the overlying frozen layer causes frost weathering on the shore, particularly where the bedrock has many fissures or joints that are widened by freeze and thaw processes. Frost weathering produces loose material for the waves to mobilise and use to erode the shore (Nansen 1922). Although periglacial processes and wave action are dominant processes now, coastal landforms have features inherited from earlier phases, such as the plateaux and escarpments produced by glacial and pre-glacial denudation. The northern, central, southern and western parts of Spitsbergen have wide, open valleys between hills and more or less rounded mountains. There are five main types of coastal landform in Svalbard: strandflats, fjords, cliffs, composite cliffs, and concave slopes (> Fig. 8.1.1.2). Reference should be made to the

map of glacial geology and geomorphology of Svalbard by Kristiansen and Sollid (1986). The term ‘strandflat’ was introduced into the Norwegian language by Reusch in 1894. In his 1912 expedition Nansen described the strandflat along the coast of  Spitsbergen, and later he advocated frost weathering and wave action as the processes responsible for the ­development of the strandflats. There are strandflats along the north coast of Nordaustlandet, on the islands of Hin­ lopenstretet and on the islands of Kongsøya, Svenskøya, Barentsøya and Edgeøya (Klemsdal 1989). The islands of Hinlopenstretet, Kongsøya, Svenskøya, Barentsøya and Edgeøya are situated east of Spitsbergen and consist mainly of Triassic sedimentary rocks with some Carboniferous and Permian formations. The strandflat on these islands may be the result long-continued post-Triassic denudation in a warm and humid climate. The denudation formed a gently, undulating topography with low relief in the eastern parts of Svalbard and on the surrounding Barents Sea floor, including a coastal plain which is the strandflat. Geo­morphological activity during Quaternary, especially glacial erosion, was of minor importance in this area and did not destroy the preglacial landforms. The fjord coasts of northern and western Spitsbergen and Nordaustlandet also had their origin in preglacial topography, which was here an upland in contrast with the lower relief further east. Long-continued weathering in a warm climate, accompanied by mass movements, shaped rounded mountains and hills dissected by river valleys. Zones of weakness in the bedrock guided the incision of valleys. Some preglacial valleys in central Spitsbergen developed an open form with a steep upper slope on the valley sides and gently declining slopes towards a broad valley floor. In Quaternary times valley glaciers at the beginning and end of the glacial periods and ice currents during maximum ice cover, further deepened these valleys. The short valleys along the west coast were transformed by valley glaciers into U-shaped valleys, and ice currents may have overdeepened the larger Isfjorden during maximum extent of glaciation. When the glaciers retreated and the sea entered the lowest parts of the valleys the fjords were formed. The slopes along a fjord may fall steeply into the sea, as in Magdalenafjord or the slope may change from a steep upper slope to a gentler lower slope down to the shore, as on the western and eastern sides of Billefjord (> Fig. 8.1.1.3). Cliffs are vertical or steep coastal slopes cut in bedrock. In Svalbard there are high sea cliffs and the low sea cliffs (> Fig.  8.1.1.2). Some cliffs are steep formerly glaciated slopes bordering fjords (> Fig. 8.1.1.3), but on the open coast waves attack the cliff base and frost weathering and mass

Svalbard and Jan Mayen

⊡⊡ Fig. 8.1.1.2 The coastal geomorphology of Svalbard.

8.1.1

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Coastal geomorphology of the Svalbard Archipelago

Glacier Fjord coast Strandflat coast High cliff coast Low cliff coast Concave coast Combined coast Glacierfront at sea 100 km

movement keep the cliff vertical or steep (> Fig. 8.1.1.4). Often there is an abrasion shore platform in bedrock or an intertidal gravelly area in front of the cliff, but where the land has been raised by postglacial isostatic rebound the cliffs emerge and become degraded to bluffs by frost weathering and mass movement. They are fronted by emerged shore platforms and shingle beaches. Such fossil cliffs are found on the northwest coast of Spitsbergen and the north coast of Nordaustlandet, all in the Hecla Hoek Complex. On some steep coasts a slope descends from a plateau and is undercut by a vertical cliff, which may be up to 30 m high (> Fig.  8.1.1.4). Such slope-over-wall profiles are found in the sedimentary bedrock of island of Hopen, at Edgeøya, Barentsøya, and on parts of the southwest coast of Nordaustlandet. Some steep coasts have a concave upper slope which descends to a gently declining concave slope towards the sea. This gently declining slope may meet the sea either in a low sea cliff, up to 20 m high or on a shore that slopes down into the sea. Landforms of this type are found along the east coast of Spitsbergen, mostly in sedimentary rocks younger than Carboniferous and along many of the fjords. In addition to these five coastal types there are ice coasts where glaciers reach the sea, often forming a vertical ice wall (> Fig. 8.1.1.5).

The shores of Svalbard are either rocky, beneath cliffs or coastal slopes, or beaches of shingle and boulders. Sandy beaches are few, but some of the spits sealing lagoons are composed of sand and shingle. Along the coast are many pocket beaches, some sandy beaches others shingle beaches (> Fig. 8.1.1.6). Beach ridges consist of pebbles and cobbles of different sizes (> Fig. 8.1.1.4) and are still forming on prograding shores. There are also fossil beach ridges that developed during Postglacial time as the land emerged from the sea due to the isostatic rebound (> Fig. 8.1.1.7). Coastal lagoons are of varying size, close to and separated from the sea by a beach ridge barrier or spits of sand and shingle. Examples are Brandalspynten by Ny-Ålesund (> Fig.  8.1.1.8) and Sarstangen, an elongated, sharply pointed arrow spit on the west side of the Strait of Forlandsundet between the islands of Spitsbergen and Prins Karls Forland. Most of the lagoons are found on the west and north coast of Spitsbergen, but lagoons are also present alongside fjords and on the islands of Barentsøya and Edgeøya (> Fig. 8.1.1.7) (Klemsdal 1986). Deltaic fans of fluvial material have been deposited near the coast or protruding into the sea along the coast of Svalbard. There are many small deltas, which show triangular fans extending out into the sea.

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Svalbard and Jan Mayen

⊡⊡ Fig. 8.1.1.3 Billefjorden, a tributary fjord to Isfjorden in central Spitsbergen. The cliff bordering the fjord has a basal talus descending into the sea.

⊡⊡ Fig. 8.1.1.4 Slope-over-wall coast on the west coast of Edgeøya Island.

2.  Bjørnøya The bedrock of the coast of Bjørnøya consists of a small area of the Hecla Hoek Complex in the southernmost part, with Carboniferous and Permian sedimentary rocks in the rest of the island and only a narrow band of Devonian formations. Small areas of Triassic rocks are found on the highest parts of the island at Miseryfjellet Mountain and Hamberg – Fuglefjellet Mountain in the south. These sedimentary rocks are gently folded and lie beneath a gently undulating plain 20–35 m above sea level.

Frost weathering, mass movement and the action of the waves and sea ice have formed a low cliff along most of the coast, as on the north coast at Bjørnøya Radio, where the cliff is between 10 and 20 m high and there is an abrasion platform in front of the headlands. Along most of the coast the abrasion platform is submerged. Headlands and bays are bordered by low cliffs, and in a few of the bays there is a shingle beach. Otherwise the shore is rocky with features produced by wave abrasion and locally ice scouring. There is a high cliff (more than 400 m) around Hambergfjellet Mountain in the southwest and the coast

Svalbard and Jan Mayen

8.1.1

⊡⊡ Fig. 8.1.1.5 The Freemanbreen glacier on Barentsøya Island with an ice front on the Sound of the Freemansundet.

⊡⊡ Fig. 8.1.1.6 Postglacial beach ridges on tombolo island northwest of Nordaustlandet Island.

of the southern tip of the island. The southeastern coast from Hambergfjellet Mountain and Fuglfjellet Mountain in the south to Miseryfjellet Mountain has a low cliff coast. The eastern coast along the Miseryfjellet Mountain has a slope-over-wall profile with is a low basal cliff, which is under attack by the waves and an upper slope formed by frost weathering and mass movement. Cliffs have been cut into the tow of a large landslide.

3.  Jan Mayen Island Jan Mayen Island is situated between 8° and 9° W and 71° N, 500 km east of Greenland and 550 km northeast of Iceland (> Fig. 8.1.1.1). The island is 53 km long and covers an area of 373 sq. km. The highest point is Mount Beerenberg, a Quaternary volcano rising 2,277 m above sea level. The island has an arctic, maritime climate with

585

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Svalbard and Jan Mayen

⊡⊡ Fig. 8.1.1.7 Location of lagoons, beach ridges, deltas and fans. 4

5

1 3

2

Photos 1 - 5 Lagoons Delta or fan Beach ridges 100 km

⊡⊡ Fig. 8.1.1.8 Delta of Bayelven River and the spit and lagoon at Brandalspynten, Kongsfjorden northwest of Ny-Ålesund.

8.1.1

Svalbard and Jan Mayen

mean monthly temperature of 5°C in August and −6°C in February and March. Average annual precipitation at the meteorological station is about 700 mm, but there are large variations with height above sea level. There are many day with fog and very few clear days. On the northwest coast of the island the cold East-Greenland Ocean current flows SW, while on the southeastern coast the warmer North-Atlantic Ocean current travels northward. Drifting sea ice normally surrounds the island from February to April, but is more persistent in the north. As the ocean currents meet around the island, air masses and air currents also converge, producing strong winds and a high wave energy along the coast of the island. From Iceland the North-Atlantic Midoceanic Ridge stretches NE as the Jan Mayen Ridge. The ridges rises from ocean depths of 1,500–2,000 m and north of Jan Mayen is a fracture zone, running NW–SE, where the sea floor is spreading about 2 cm a year. The ridge started to form on the sea floor in the Eocene, but the oldest rocks above sea level only go back 5,00,000 years. The island is thus relatively young, composed of basaltic lavas and tephra. There are many minor craters on the southwestern part of the

island, but the Beerenberg volcano dominates the island, the lower part (up to 1,500 m above sea level) a shield volcano with a slope of 15°, while the upper part is a stratovolcano with slopes of 30°. In September 1970 there was a 3 week eruption on the lower northeastern flank of the mountain. Lava flowed out from craters along a 5 km long fissure zone and down towards the sea to add an area of 4 sq. km of new land. During the following winter more than half of this land was eroded away by waves. Further eruptions in 1985 produced salients of lava extending into the sea. There are many days with stormy weather, and Jan Mayen is exposed to strong wave action from all directions: from the North Atlantic, the Greenland Sea and the Arctic Ocean. Generally basaltic lava is resistant towards erosion while volcanic gravel and ash are more easily eroded by the waves, transported along the coast and accumulated as sand and shingle beaches. There are high sea cliffs (10 m to more than 200 m), low cliffs ( Fig. 8.1.1.9). The longest continuous high sea cliffs are

Coastal geomorphology of Jan Mayen

8° west Greenwich Lava 1985 Kokssletta Lava 1970

71°North

Engelskbukta

8° west Greenwich

Kvalvossbukta

Rekvedbukta

Sea stack Rocky shore

Guineabukta

1 km

Greenland

80°N

Svalbard 30°E

Beach ridge Glacierfront in sea Low sea cliff

Glacier

70°N

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Jan Mayen 0°

Slide fan Ny lava

10°E

10°W

High sea cliff 9°west Greenwich

20°E

20°W

Iceland Polarcircle

way

Sandy / stoney beach

Nor

⊡⊡ Fig. 8.1.1.9 Coastal geomorphology of Jan Mayen Island.

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Svalbard and Jan Mayen

south and east of Beerenberg and on the SE and SW coasts. Cliff erosion has produced stacks, arches, caves and notches. In front of the high cliffs there is either a rocky shore where wave abrasion has cut an abrasion shore platform or a stony beach covering this platform. Extensive low cliffs fronted by rocky shores are found in the southwest be­­tween Kraterflya and Guineabukta and along the northernmost coast of the island. On the west coast (west of Kronprins Olavs bre), the land surface slopes gently into the sea, by way of a rocky shore or abrasion platform. At Kokssletta a 4,000 year old glacial moraine extending on to an emerged (6–10 m) shore platform is overlain by a more recent lava flow. There are alternations of rocky shores and shingle beaches in front of these cliffs. Beaches are derived from cliffs of weathered lava and volcanic sand and gravel by frost action, slumping and wave scour. Shingle beach ridges have formed in the bays of Rekvedbukta and Eng­ elskbukta, in places enclosing lagoons, as at Sørlaguna, which dries out during the summer, and Nordlaguna,

which is 40 m deep. A barrier beach 11 km long encloses Lagunevallen on the south coast, and a similar barrier fronts Stajonsbukta on the north coast, where there is a sandy beach north of Sørlaguna. Emerged beaches occur in several places along the coast of Jan Mayen.

References Klemsdal T (1986) Lagoons along the coast of the Svalbard archipelago and the island of Jan Mayen. Norsk Geogr Tidsskr 40:37–44 Klemsdal T (1989) Landforms – Svalbard and Jan Mayen, 1:1,000,000. Map. Norwegian National Atlas. Hovedtema 2: Lanformer, berggsunn og løsmasser. Kartblad 2.1.3 Kristiansen KJ, Sollid JL (1986) Svalbard map of glacial geology and geomorphology, 1:1,000,000. Norwegian National Atlas. Statens kartverk, Hønefoss, Norway Nansen F (1922) The Strandflat and Isostasy. Videnskapsakademiets Skrifter. I Matematisk Naturvitenskapelig klasse nr 11:313 pp Winsnes TS (1988) Bedrock of Svalbard and Jan Mayen. Geological map 1:1,000,000. Norsk Polarinstitutt. Temakart nr 3:12 pp

8.2 Sweden

John Norrman*

Introduction Sweden (>Fig.  8.2.1) forms the Archean basement marginal to the Baltic-Russian sedimentary basin to the east, the Danish-German basins in the south, and the Caledonian mountain range in the west. Within the Swedish shield area are a rich diversity of gneisses and granites. In the Baltic Sea is a submerged scarpland of lower Palaeozoic rocks, in which the main features are the northwest margins of the Ordovician and Silurian limestones which make up the islands of Oland and Gotland. The main part of the Bothnian Sea, excluding the Bothnian Bay, is also floored by Palaeozoic sediment, but there are only a few scattered outcrops on the coast. During several stages of the Quaternary, Sweden and the area now occupied by surrounding seas were completely glaciated. The final glaciation was during as Weichselian substage, which culminated only some 30,000 years ago. When the front of the receding ice reached Scania, the southernmost province of Sweden, about 14,000 years ago, a complicated history of coastal evolution began, dependent on the rate of deglaciation, land uplift by isostatic rebound, sea level variations and shifting outlets from the Baltic basin. The altitude of the Holocene marine limit, the highest postglacial metachronous coastline (>Fig. 8.2.1), rises from 20 m in southern Scania to about 220 m at the southern end of the Bothnian Sea and a maximum of 280–285 m in the Hoga kusten area at latitude 63° N. From there it falls to 160–220 m within a broad uplifted archipelago at the northern end of the Bothnian Bay. Along the west coast the highest Holocene coastline rises from 50 m in the northwestern part of Scania to about 170 m at the Norwe­gian border. The inland ice produced a till cover of variable texture and thickness and glaciofluvial sediment of two distinctive types: coarser sands that were deposited in ice tunnels and meltwater channels to form eskers and deltas, and suspended silt and clay deposited below sea level and in ­ice-dammed lakes. The suspended load formed density currents that annually distributed varved sediment to the lowest areas. The average thickness of the till is not more

than 3–4 m, which means wave action exposed the ­glacially polished bedrock on coastal slopes. Continuous reworking of material already washed down in the swash and breaker zone effectively separated the suspended matter, which settled on top of the varved sediment. In this way and by man’s exploitation, a landscape with forested hills, scattered bedrock outcrops, and cultivated sediment plains and valleys came into existence. The present rate of uplift varies from zero in southern Scania to 0.4–0.5 m/century in the Stockholm area and a maximum of 0.9 m/century on the coast of the Bothnian Bay. The extent of uplift has been measured with reference to former sea level bench marks cut in rocks at Oregrund and Gefle, and to the inscription carved by Charles Lyell on 18 July 1834 at Gullhomen. The position of the highest Holocene coastline and the rate of uplift mean that large areas have in comparatively recent time been operated on by coastal processes, but for such a short duration at each level that the coastal influence is often only detectable by a sorting of the topsoil. At strongly exposed sites or at levels where the emergence slowed down or even reversed to a transgression, well-developed relict shores may be found (>Fig. 8.2.2). They are especially well preserved where a transgression has been followed by a rapidly falling lake or sea level, as for example in the lowering of the Baltic Ice Lake to sea level about 10,300 years ago. Bedrock morphology; soil characteristics, and the history of land uplift imply that the present coasts are generally in an early stage of development. The inherited bedrock morphology is, in detail, still characterised by glacial sculpture, but on the macro scale it is often far older, as more or less modified Pre-Cambrian and Mesozoic peneplains dominate large coastal environments. With regard to the generally primary character of the Swedish coast a regional division based mainly on the inherited pre-Quaternary macromorphology has been chosen for this review. In the archipelago coast of the Bothnian Bay, large scale bedrock elements with a NW-SE orientation dominate the morphology. The same orientation is found in the valleys of large rivers, and their ­preglacial continuations can be traced across the Bay

*Edited version of chapter 8.2 (Sweden) in The World’s Coasts: Online (2003) by John Norrman (University of Uppsala, Sweden). All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.2    

Sweden

⊡⊡ Fig. 8.2.1 The coastal regions of Sweden. KEY: (A):- The archi­ pelago coast of the Bothnian Bay (B):- The coastal plain of the Northern Quark (C):- Höga kusten – the high coast of the Bothnian Sea (D):- The coastal plain of the Bothnian Sea (E):- The rocky barrier coast of the Southern Quark (F):- The archipelago coasts of the Baltic Sea: (F1):- The Stockholm archi­ pelago (F2):- The Södermanland ­archipelago (F3):- The östergötland archipelago (F4):- The Blekinge archipelago (G):- The coastal plain of the Kalmar sound (on the mainland) (H):- The cuesta coasts of the islands of öland and Gotland (I):- The open horst and glacial drift coasts of Skåne (Scania) (J):- The lowland coast of the Kattegat (K):- The archipelago coast of the Skagerrak. Arctic circle

Stockholm Peneplain Dissected Peneplain

BA

Y

AN

D

A

E BOTHNIA TH N

FINLAND

SE A

B

C

D

N 60° eM

älar

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S

TIC SEA

Lak

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an

ia

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E2

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tlan

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lan

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Ö

590

to the Finnish coast. At the margin of the Bay this ­downwarped, presumably Tertiary, landscape forms a riatype archipelago characterised by open basins bordered by rather large islands and submarine sills. The same type of coast is found in the province of Blekinge in south-eastern Sweden (Lidmar-Bergström 1982). In the Bothnian Bay glaciofluvial eskers run along the preglacial valleys and ice-front meltwater accumulations and series of small end moraines add oriented elements to the bedrock morphology. The coast is generally low. Slope measurements by Håkanson (1982) together with present land uplift (about 9 mm/year) indicate that a 1.1 m broad strip of land has been added annually. The coastal plains of the Northern Quark, the Bothnian Sea and the Baltic Sea are fringed with islets and skerries, but lack the width of the archipelago coasts. These low coasts exhibit Pre-Cambrian peneplain surface fragments which gently dip under Palaeozoic sediment in the Both­ nian Sea and the Baltic Sea. The present coastline is largely dominated by wave-washed till in the more exposed parts and by re-deposited sediment and gyttja formed beneath a wide zone of sedges in shallow sheltered bays. Because of the continuous land uplift, fine material deposited below the limit of wave action can be re-suspended later. It has been estimated that most of the annual sedimentation off the Swedish east coast emanates from re-suspension. Inland the peneplain that forms the basis of the Bothnian coastal plain is an undulating landscape that rises in steps, especially well illustrated along the river valleys, and rapidly reaches relative heights of more than 100 m. This morphology meets the coast at the northern end of the coastal plain in a region named Höga kusten (the high coast of the Bothnian Sea), where the rocky hills reach heights of more than 300 m and where the maximum depth off the coast is 250 m. On this coast the ­highest postglacial coastline in Sweden stands at 285 m. Most of the glacial soil has been stripped by wave action from the exposed steep hillsides to form leeward delta-like shore deposits, with only very coarse boulder terraces on the seaward side (>Fig. 8.2.3). During the land uplift the glacial deposits in the valleys have been repeatedly reworked at lower levels. This material now forms terraces along the rivers and deltas in present-day estuaries. At the Southern Quark the wide, almost horizontal Pre-Cambrian peneplain comes to an abrupt end along a rocky coast with major joints oriented parallel to the coastline. The maximum offshore depth is 200 m. The broad Stockholm archipelago coast is dominated by glacially abraded bedrock structures and fissure valleys etched in the same peneplain (>Fig.  8.2.4). South of this region the archipelago coasts are characterised by

Sweden

    8.2

⊡⊡ Fig. 8.2.2 Raised beaches and dunes along the shore of Haparanda Sandskär, 20 km west of the Finnish border on the north coast of Bothnian Bay.

⊡⊡ Fig. 8.2.3 Boulder berms at 230 m above sea level on the exposed eastern slope of Högklinten in Mjällom, Höga Kusten.

E-W faults which cut off N-S and NW-SE fissure valley systems (Rudberg 1970). The Swedish part of the central Baltic Sea is occupied by a submerged scarp of Lower Palaeozoic sedimentary rocks, in which the main features are the north-western margins of Ordovician and Silurian limestones which dip towards the east and form cuesta coasts on the islands of öland and Gotland. The western cliff coast of öland (>Fig. 8.2.5) runs along an escarpment in Ordovician rocks that can be followed in a curved line across the Baltic Sea to the ­northern coast of Estonia. Almost parallel to that is an escarpment in

Silurian rocks which forms a steep submarine slope along the north-western coast of Gotland. At this coast frost action and wave action have alternately cut vertical cliffs and steep slopes with uplifted shingle berms faced by abrasion platforms (Rudberg 1967). According to more recent studies the abrasion platforms, now in existence, did not start to form until about 5,500 years ago (>Fig. 8.2.6). Having been covered by the Late Cretaceous Sea, the  province of Skåne (Scania) belongs geologically to the European continent rather than to Fennoscandia. The Archean basement was broken up, mainly in the Permian,

591

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8.2    

Sweden

⊡⊡ Fig. 8.2.4 In the outer part of the Stockholm archipelago, swarms of small skerries with glacially sculptured rocks are common. In this view from Kallskär, where the inland ice moved from the left, stoss-and-lee morphology is well illustrated.

⊡⊡ Fig. 8.2.5 Active cliff fall of limestone directly on to the abrasion platform on the west coast of öland. A beach of longshore drifted shingle in the background.

into horsts and grabens along NW-SE fracture zones. The horsts are illustrated by the peninsulas of Kullen protruding into the sea on the west coast, and the grabens by the sandy bay between these, as well the large Hanö Bay to the east. On the horsts are steep fractured cliffs and emerged Holocene shore features (>Fig. 8.2.7). Except for the horst areas, glacial drift deposits are up to 40–50 m thick in Skåne, and although low coasts ­predominate there are active cliffs cut in till and glaciofluvial sediment along elevated sectors. The isoline of present zero uplift runs in NW-SE through central Skåne, which together with the abundance of easily eroded soils explains

why Skåne is the only mainland province with notable coastal erosion. On the Falsterbo peninsula there are sub-parallel dune ridges that formed by accretion of wind-blown sand on successively built longshore spits. The origin of the roughly 20 km broad lowland bordering the Kattegat in the province of Halland has been compared with the Norwegian strandflat, and certainly seems to have had a complex history. The gross mor­ phology of the bedrock surface was initiated by deep weathering in the Mesozoic, and an uneven erosion surface formedin the Early Cretaceous was buried in the Late

Sweden

    8.2

⊡⊡ Fig. 8.2.6 Shore stacks on a raised shingle beach at Fårö on the north-western coast of Gotland, developed in soft Upper Silurian marlstone. The stacks originate from pillarshaped resistant concretions in the limestone. White shingle marks the upper limit of present swash action.

⊡⊡ Fig. 8.2.7 Shingle berms built successively from 9 m above present sea level, starting about 6,500 years ago, on the north coast of the Bjäre horst at the southern end of the Kattegat coast.

Cretaceous, then exhumed and remodelled in the Neogene (­Lidmar-Bergström 1982). Especially in the southern part,  coastal landforms are largely the product of recent shore processes.  Large open sandy bays are backed by broad dune belts (>Fig.  8.2.8) comprising active backshore dunes and an inner zone stabilised by pines mainly planted during the ­nineteenth century. Erosion began with the start of agricultural activity as early as the sixteenth century, and during the last 50 years the attraction of sandy beaches for recreational activities has rapidly

increased and caused environmental problems, notably destruction of dune grasses, which initiates wind erosion (Norrman et al. 1974). The archipelago coast of the Skagerrak in the province of Bohuslän owes its general character to a strongly dissected Pre-Cambrian peneplain in which valleys were deepened by glacial excavation to form fjords. Much of the bedrock is exposed because the thin till cover was washed down into the valleys by wave action as the land was uplifted.

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8.2    

Sweden

⊡⊡ Fig. 8.2.8 Typical dune zonation in Tylö Bay on the low southern coast of the Kattegat. A wide sandy beach is backed by grass- covered dunes with many trampled tracks, and then a broad belt of planted pine trees.

References Hakanson L (1982) Coastal morphometry the coast of the Gulf of Bothnia. In: Willer K (ed) Coastal research in the Gulf of Bothnia. Biol Mono, 45:9–33 Lidmar-Bergstrom K (1982) Pre-Quaternary geomorphological evolution in Southern Fennoscandia. Sveriges Geol Unders Ser C:785

Norrman JO (1992) Coast and Shore. Sea and Coast. The National Atlas of Sweden. SNA Publisher, Stockholm Norrman JO et al (1974) Investigations of dune morphology in southern Halland (in Swedish). Statens Naturvårdsverk PM 500, Stockholm Rudberg S (1967) The cliff coast of Gotland and the rate of cliff retreat. Geogr Ann 49A:283–298 Rudberg S (1970) Geomorphology. Atlas over Sverige, 5–6

8.3 Finland

Jouko Alestalo · Olavi Granö  

This chapter is based on a more detailed article written in  cooperation with the late Dr. Jouko Alestalo of the University of Oulu, Finland, which appeared in the first edition of this volume.

1.  Introduction The coast of Finland is a mosaic of headlands, islands, bays and straits. There are about 73,000 islands, of which 52,000 are less than 1 ha in area and only 15 exceed 50 km2. Measured on the 1:10,000 map the coastline is 47,518 km long (Granö 2007), varying in width from about 120 km in the southwest (the Åland archipelago) to less than 5 km along the Gulf of Bothnia. About 41.9% of the shoreline is rocky, 41.6% is composed of glacial drift, 10.4% silt, clay and marshes, 4.8% sand and gravel, and 1.3% artificial (Granö et al. 1999). The islands are of three kinds: rocky, especially in the south and southwest, where they are often encircled by glacial drift partly covered with silt and clay; morainic on the coast of the Gulf of Bothnia and eskers, with sand and gravel partly covered by silt and clay. These are often in

sequence, the outer islands being rocky or boulder moraine and those further landward having a mantle of glacial drift, overlain in the innermost zone by a horizon of silt and clay. Esker sands and gravels are overlain in places by an emerged clay layer, which is an inheritance from the Litorina Sea. Cliffs are rare, usually cut in glaciofluvial deposits, but such erosion is countered by continuing emergence. Steep rocky coastal slopes have generally been shaped by prior glaciation and have been modified very little, if at all, by marine erosion. The morphology of the coastline is largely determined by structure of the Pre-Cambrian bedrock, as revealed by selective erosion, especially during glaciations, and has been shaped to a lesser extent by the glacial drift superimposed on this bedrock. In the south–west erosion of cliffs in soft formations such as eskers has been fairly rapid on coasts exposed to strong wave action, because of the long ice-free period each year and the relatively slow rate of land uplift. Elsewhere on the outer islands stony outwash shores are dominant. Land uplift has led to the creation of successive beach ridges in places on the sand and gravel shores (>Fig. 8.3.1). A well-developed wave-shaped morphology is found on

⊡⊡ Fig. 8.3.1 Beach ridges on Hailuoto, Bothnian Bay. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8. 3

Finland

⊡⊡ Fig. 8.3.2 Long sandy beach at Lohtaja, backed by partly vegetated dunes. Emergence and sand deposition have cut off a former inlet to create a lake. (Courtesy Hannu Vallas.)

⊡⊡ Fig. 8.3.3 Pattern of land and water, related to structure in granitic bedrock, SW Finland. (Courtesy Geostudies.)

Finland

the coast of the Bothnian Bay, in spite of the rapid rate of land uplift, because the shores are exposed to strong wave action and the surface deposits are largely composed of sand. There are sandy beaches backed by dunes (Hellemaa 1998), often with irregular outlines (>Fig. 8.3.2). The Åland archipelago (>Fig. 8.3.3) has 75% rocky shores, the proportion of glacial drift and silt and clay increasing landward and towards the southwest coast, where no more than a half of the shores are rocky, 30% formed by glacial drift and 11% by clay and silt. These proportions are similar in the western part of the southern (Gulf of Finland) coast, while the eastern part has only 27% rocky shores but 46% composed of glacial drift and 12% of sand and gravel. On the Bothnian Sea coast north of 61° N the shores are 80% glacial drift and only 6% rocky. The rivers deposit large quantities of silt, sand and gravel in their mouths, which together with glaciofluvial material, accounts for 93% of the Bothnian Bay shores (Granö and Roto 1989). The climate of Finland is dominated by southwesterly winds associated with the passage of depressions. The summers are mild but short, and it is sufficiently cold in winter for ice to form on the shores. Helsinki has a mean monthly temperature of −6.2°C in January, rising to 17.2°C in July, and an average annual rainfall of 627 mm, a third of which falls as snow. Coastal waters and shores around the islands are covered by stationary ice for as long as 5.5 months on average at the head of the Bothnian Bay and over 3 months on the south coast. Low-pressure and winds cause water levels to rise and detach this ice from the shores, leading to rafting, piling up of the ice on the shores and ice-push depositing stony material on the backshore. The shores of bays and narrow straits are also shaped by thermal ice movement (Alestalo and Häikiö 1976). Wave action, generated mainly by the southwesterly winds, is related to the length of the fetch. Significant wave heights on the southern coast are 1.9–2.2 m, with occasional waves of up to 7 m, while those in the southwest are over 3 m, with occasional waves of up to 14 m, those in the Bothnian Sea 2.4 m those and in the Bothnian Bay less than 2 m. Storms associated with cyclones moving on the polar front are found mainly in autumn, when the shores of the Gulf of Finland and the south western islands are scoured by waves whipped up by the southerly winds. The wind shifts to the west in the latter part of the cyclone, when the waves wash the shore of the Gulf of Bothnia. In spring the disruption of shore ice by waves is accompanied by the plucking of rocky outcrops, contributing to exfoliation (>Fig. 8.3.4).

8. 3

⊡⊡ Fig. 8.3.4 Exfoliation on a granite shore, Seili, SW Finland. (Courtesy Geostudies.)

Tidal fluctuations in the Baltic Sea are of the order of only a few centimetres, but winds and changes in air pressure can give rise to major variations in water level at the heads of the Gulf of Finland and Gulf of Bothnia, ­especially in early winter. Sea level can rise more than 2 m at the head of the Gulf of Bothnia, 1.5 m at the eastern border of Finland, and about 1 m in southwest Finland. There are weak coastal currents running northward in the Gulf of Bothnia and westward in the Gulf of Finland. Sea temperatures rise to a summer maximum of 16°C–18°C in the Gulf of Finland and 14°C–16°C in the Bothnian Bay, the salinity of the Baltic Sea being about 0.6% off the southwest coast and diminishing to 0.1% at the heads of the gulfs. The estuaries are almost fresh water, but show ­gradients in sedimentation. Reeds and rushes are extensive on sheltered shores (>Fig. 8.3.5). Nearshore plant growth includes Cladophora spp., with Fucus spp. in the more brackish areas.

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⊡⊡ Fig. 8.3.5 Sedges and reeds on the NW shore of Hailuoto. (Courtesy Geostudies.)

⊡⊡ Fig. 8.3.6 Rates of land uplift around the Gulf of Bothnia, based on Finnish and Swedish geodetic surveys. (Courtesy Geostudies.)

Finland

⊡⊡ Fig. 8.3.7 Historical changes on the coast of the Bothnian Bay near Nykarleby between 1773 and 1967 as a result of isostatic land uplift. (Courtesy Geostudies.)

8. 3

Russian border. This uplift has been partly offset by a eustatic sea level rise of about 1 mm/year. The eustatic rise has increased since 1975, and is now greater than the rate of land uplift in the eastern Gulf of Finland. This glacio-­ isostatic land uplift has caused the emergence of hundreds of metres of land per century on the gently sloping shores of the Bothnian Bay, cutting short the development of littoral landforms. As emergence proceeds the mouths of some rivers become rocky rapids, and the smaller rivers become isolated as lakes. Harbours become shallower, requiring dredging, and some have been stranded in inland positions. It has been estimated that the increase in land area in Finland amounts to 700 sq. km every hundred years as a result of land uplift. If, to this, are added the effect of sedimentation  and vegetation growth, the figure can be raised to 1,000 sq. km (Jones 1977).

References

The whole coastline of Finland has been affected by land uplift (>Fig. 8.3.6), which amounts to 9 mm/year at its most rapid, on the coast of the Bothnian Bay between Oulu and Vaasa (>Fig. 8.3.7), and is least pronounced about 3 mm/year on the Gulf of Finland shore close to the

Alestalo J, Häikiö J (1976) Ice features and ice-thrust shore forms at Luodonselkä, Gulf of Bothnia, in winter 1972/73. Fennia 144:5–24 Granö O (2007) Coastal studies in Finland. Yearbook of the Estonian Geographical Society 36:156–175 Granö O, Roto M (1989) Zonality in the Finnish coastal environment. Essener Geographischen Arbeiten 18:269–281 Granö O, Roto M, Laurila L (1999) Environment and land use in the shore zone of the coast of Finland. Publicationes Instituti Geographici Universitatis Turkuensis 160:76 Hellemaa P (1998) The development of coastal dunes and their vegetation in Finland. Fennia 176:1–245 Jones M (1977) Finland. Daughter of the sea. Dawson, Folkestone, Kent

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8.3.1 Russian Gulf of Finland

Alexey Porotov

1.  Introduction The Gulf of Finland occupies a large tectonic depression between the crystalline shield of Fennoscandia and the East European platform. It is now relatively stable, with recent neotectonic movements of less than ±1 mm/year. The crystalline basement outcrops only in the northwestern and south-west part of the gulf. In the northwest it forms the numerous small islands and peninsulas of  Vyborg Gulf. The southern coastline has cliffs (klint or  glint) up to 30 m high, cut in carbonaceous rock. Much of the coast of the Gulf of Finland has a cover of Quaternary deposits including glacial till and fluvial, marine and aeolian sands of varying thickness. The retreat of the sea following the Litorina phase has resulted  in terraced slopes with abrasion benches. According to Markov (1934) these are not necessarily evidence of ­stillstands, and could have formed during a steady emergence. The Gulf of Finland is exposed to the prevailing westerly winds and storms, which produce waves and currents that drift sediment alongshore into the inner part of the Gulf. The absence of significant fluvial sediment input contributes to instability of the coastline, which is receding at an average rate up to 1.5 m/year. Narrowing eastward, the Gulf of Finland amplifies storm surges that can exceed 4 m above mean sea level. There have been episodes of coastline recession of up to 5 m on low-lying sectors. Storm surge flooding causes extensive damage in coastal areas, particularly in the highly populated zone around St Petersburg. There is flooding of low-lying coastal areas, the zone subject to submergence reaching 5 km at the head of the bay near west-facing Sestroretsk City. In winter there is extensive ice in the Gulf of Finland. This impedes wave action, but when it breaks up in spring it is jostled by waves, forming push ridges on gravelly beaches and scouring the base of cliffs (Bachmanov 1935).

Erosion of the coastline in the eastern part of the Gulf of Finland has been accelerated by human activities such as dredging of sand and breakwater construction in many places.

2.  The Gulf of Finland Coastline East of the Finland border is the large funnel-shaped estuary of the Vyborgskiy Zaliv. Beyond this the north coast of the Finnish Gulf has cliffs up to 5 m high cut into glacial drift deposits. These are fringed by beaches of sand and shingle up to 10 m wide, derived mainly from the glacial drift (> Fig. 8.3.1.1). The rate of cliff recession increases from 0.3 m/year where the glacial deposits contain a large proportion of gravel) to 2 m/year where the glacial till is sandy. Apart from these few sectors of cliff cut in Pleistocene glacial drift the coastline is dominated by sandy beaches backed by a wide zone of Holocene sandy beach ridges capped by dunes. Many beaches are derived from outcrops of Pleistocene glacial drift. The width of the beach depends on the type and abundance of sediment. Coasts with little sediment have narrow (15–20 m) beaches which consist of sand and pebbles, the pebble gravels impeding active coastal recession (> Fig. 8.3.1.2). Groynes built in the 1930s had little effect on these beaches. Broader beaches (up to 100  m wide) occur on coasts receiving longshore drift, as at the head of Sestroretsk Bay, or updrift of the large Zelenogradsk breakwater. The beach is narrower and eroding downdrift of this break­ water (> Fig. 8.3.1.3). At the head of the Gulf is the large port city of St Petersburg (formerly Leningrad) at the mouth of the Neva River, which carries the outflow from Lake Ladoga. The southern coast of the Russian Gulf of Finland is generally low-lying, and consists of a series of broad promontories with some cliffed sectors, separated by bays with sandy beaches, along to the Estonian border.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.3.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Russian Gulf of Finland

⊡⊡ Fig. 8.3.1.1 Cliff cut in glacial boulder loams on the north coast of the Finnish Gulf near Repino.

⊡⊡ Fig. 8.3.1.2 Boulder groyne on the north coast of the Finnish Gulf.

Russian Gulf of Finland

8.3.1

⊡⊡ Fig. 8.3.1.3 Eroding sand beach nearly 30 m wide downdrift of the Zelenogradsk breakwater.

References Bachmanov BM (1935) Springtime accumulations of ice. Priroda, Moscow, 8

Markov KK (1934) Criteria of transgression and regression. Proceedings of the First All-Union Geographic Congress 3

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8.4 Estonia Kaarel Orviku · Are Kont · Hannes Tõnisson

1.  Introduction

The cliffed (klint) north-eastern coast around Ontika has been straightened by erosion, whereas the beach-fringed Narva Bay has an outlined smoothed by deposition coast (Orviku and Romm 1992). Coasts straightened by a combination of erosion and deposition can be found on the northern shore of Köpu Peninsula in the western part of Hiiuma Island and around the Gulf of Livonia. Orviku and Sepp (1972) have illustrated the stages in landform evolution on the coasts of the west Estonian archipelago. Coastal evolution in Estonia has been influenced by changing sea levels: the Ancylus transgression, followed by the Litorina transgression and an ensuing regression (7,000–5,000 years ago). During the Litorina transgression there was extensive erosion, producing cliffs and shore platforms and generating large amounts of sand and gravel, with residual boulders where glacial deposits (till) had been dispersed. During the ensuing regression sand, gravel and boulders on the emerging sea floor were shaped into beaches, beach ridges and dunes by wave and wind activity.

The coastline of Estonia is about 3,790 km long, including Hiumaa, Saaremaa and over 1,500 smaller islands of the Estonian Archipelago. It includes the sectors of the coast exposed to waves generated by the prevailing westerly winds, with NW waves dominant along the north-facing segment beside the Gulf of Finland, and the sectors that are relatively sheltered on the inner coasts of islands and along the Gulf of Livonia (Riga) to the south. The islands of the west Estonian Archipelago are separated from the mainland by Väinameri (Moonsund). The classification of the Estonian coastline (>Fig. 8.4.1) is based on the concept of wave processes straightening the initially irregular outlines by erosion of capes, deposition in bays, or a combination of the two (Orviku and Granö 1992). Much of the coast (77%) is irregular, with capes and bays either in hard bedrock, as in Lahepere Bay west of Tallinn, or in unconsolidated Quaternary deposits (­notably glacial drift), as in Lahemaa National Park east of Tallinn.

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DENNAM



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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8.4

Estonia

Postglacial isostatic movements resulted in land uplift ranging from about 45 m in southern Estonia up to 75 m on the northern coast since the Ancylus Lake period. Beaches are often backed by beach ridges and dunes, the inner and older of which have been raised by the tectonic movements. Tilting has continued on either side of a zero isobase that runs SW-NE through Riga in Latvia, with land uplift of about 1 mm/year at Pärnu, 2 mm/year at Tallinn and 2.8 mm/year on the northwestern coast (>Fig. 8.4.2) (Vallner et al. 1988). Tides are negligible on the Estonian coast, but sea level is generally higher in winter than in summer. Westerly and southwesterly winds predominate producing waves from these directions on west-facing sectors of the coast, but wave energy is generally low, especially in places where near-shore is shallow and boulder-strewn. Wave activity is stronger during storms, and in storm surges the onshore winds and low barometric pressure may raise sea level nearly 3 m above the Kronstadt zero (benchmark for the eastern Baltic Sea). This is when major changes take place along the coast, with cliffs undercut at a higher level, high beach ridges formed, and large quantities of sand moved alongshore. During winter (November–April), a shore ice fringe develops and prevents wave activity, but when it breaks up in spring waves may drive it onshore, piling it up as 10–15 m high hummocks. Ice driven onshore scours sea floor sediment, displacing sand and gravel, and even boulders shoreward, damaging trees and buildings on the coast (>Fig. 8.4.3).

0

The depression of the Gulf of Finland and the North Estonian Klint are formed in the contact zone between the crystalline shield of Fennoscandia and the rest of the East European platform, and have been shaped by erosion of the pre-Quaternary large rivers (Raukas 2005). The soft Vendian and Cambrian clays and sandstones that were once exposed in the southern part of Finland and everywhere in the bottom of the Gulf of Finland, were eroded. The stronger Ordovician limestone hindered erosion, forming a high klint (Orviku 1982; Raukas 2005) and protecting the underlying softer rocks from rapid erosion. The klint margin and the terraces were also eroded by continental ice. The Baltic Ice Lake formed after recession of the glacier and the subsequent development stages of the Baltic Sea have washed the base of the klint, subjecting it to continuous slow erosion. The crystalline basement in Estonia lies at a depth of 150–500 m under the Lower Palaeozoic sedimentary rocks. Quaternary deposits of various thickness cover the Palaeozoic bedrock. Pleistocene glacial, fluvio- and limnoglacial deposits together with Holocene mineral and organic sediment form the Quaternary cover. Lowland coasts are predominant in Estonia. Higher coasts with cliffs occur where bedrock outcrops are present or glacial marginal ridges are intersected by coastline forming scarps in Quaternary deposits. Waves and currents (especially during heavy storms), long-shore drift, onshore winds, and human activity are the main agents of coastline evolution. Rivers are small

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⊡⊡ Fig. 8.4.3 Shore ice pushed tens of metres into the coastal pine forest on the northern coast of Saaremaa.

and alluvial deposition is very limited. There has been an increase in storminess in the coastal waters of Estonia with nine major storms between 1965 and 2001 with an intensity that previously occurred only once or twice in century (Orviku et al. 2003). Frequently raised sea level by storm surges, a general absence of shore ice cover with unfrozen shore sediment in milder winter conditions allow waves to attack the coast and shape the beaches even in winter. Despite tectonically uplifting coast, beach erosion ­attributable to increased storminess has become evident in Estonia in recent decades.

2.  The Estonian Coastline Near the Russian border in the northeast is Narva Bay, where a long gently curving sandy beach is backed by numerous parallel beach ridges covered with dunes representing intermittent progradation in Holocene, during the post-Litorina emergence. The sand has come partly from erosion of cliffs to the west, partly from glacial and fluvioglacial deposits on the Kurgalovo Peninsula to the north, and partly from the River Narva, the yield from which has diminished after the dam construction upstream. With a reduced sand supply, the seaward margin of these dunes is now eroded forming abrupt scarps and receding. In the 1980s the beach was re-nourished, and a breakwater was built at the mouth of the River Narva (Orviku and Romm 1992). The ice fringe, which forms in

winter is disrupted by spring storms and is piled up to 7 m high on this beach. The Baltic klint becomes higher and is located closer to the coastline west of the Narva Bay past Sillamäe to Kunda. It is cut into Cambrian sandstone, clay and shale in the lower part of the klint covered by more resistant to erosion  Ordovician limestone. The highest cliffs in average 20–30  m and rising to a maximum of about 56 m (>Fig.  8.4.4) are between Ontika and Kalvi. There are accumulations of basal rocky talus and beaches of cobbles and boulders derived from cliff erosion, and the near-shore area is a boulder-strewn subtidal abrasion platform. The klint is now only locally and episodically attacked by waves, and is generally rather stable. To the west, in Lahemaa, the coastline becomes irregular, with four large peninsulas separating deep bays. At the head of each bay is a sandy beach, backed by beach and dune ridges that rise landward as the result of the Holocene land uplift, as at Võsu. In Eru Bay, erosion of glacial deposits on the Käsmu and Pärispea peninsulas has resulted in bay-head accumulation of sand and silt, with reeds and bushes colonizing silty flats. In Lahemaa, the coast consists of low scarps cut into glacial and fluvioglacial deposits, varved clays and gravel, and the shore platforms cut into these soft sediment are strewn with erratic boulders (>Fig. 8.4.5) and fringed by reed beds. There are bay-head beaches of sand and gravel, the finer silt and clay having been withdrawn in suspension seaward and deposited on the bay floor. More promontories and bays extend to the

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⊡⊡ Fig. 8.4.4 Up to 56 m high cliff at Päite, northern coast of Estonia. The talus at the foot of the cliff is covered by permanent forest vegetation.

⊡⊡ Fig. 8.4.5 Abraded till shore on Mohni Island, Gulf of Finland. Finer-grained particles are washed out by waves and remaining boulders provide a good natural protection against further erosion.

large Viimsi Peninsula, and the capital city of Tallinn stands behind the next broad bay. Here the Ordovician cliffs overlie the Cambrian formations. In Tallinn Bay, there are 4 km of breakwaters and jetties, and in 1980 a 2.5 km long sea wall was built to enclose and reclaim shallow bay-head area for highway

c­ onstruction and the Olympic Yachting Center (Martin and Orviku 1988). An artificial beach emplaced at Pirita has become wider as the result of natural sand accretion. The northwest coast of Estonia is irregular consisting of valley-mouth bays with sandy beaches and reed beds and limestone headlands with cliffs. In Lahepere Bay,

Estonia

between the capes of Türisalu and Pakri, a wide beach has formed, backed by dunes and fronted by submerged sand bars. The sand has been derived from Lower Ordovician sandstone in the bordering cliffs. On Cape Pakri, a cliff has been cut into relatively soft Lower Ordovician sandstone, which is overlain by a harder limestone. The limestone upper cliff, which is undermined by wave-cut notches, frequently collapses forming a talus. The talus is subsequently reduced by wave activity, and turned to gravel beach. As the coast turns southward, there are two large irregular inlets, Haapsalu Bay and Matsalu Bay, opening to the sector of the Baltic Sea that is sheltered by Hiiumaa Island. Many small rocky islands occur in the northern part of Moonsund Strait. Because of continuous land uplift, cliffs are usually eroded only during heavy storms. The Matsalu Bay at the mouth of River Kasari is bordered by extensive reed bed, which has spread rapidly seaward with the emerging near-shore area. Seaweeds, such as Cladophora, Potamogeton and Chara, grow prolifically in the near-shore waters. A former island Haeska has been attached to the mainland as a result of land uplift at the mouth of the Matsalu Bay. Its formerly beach-fringed coast is now fringed by reed bed. By contrast, Kumari Island to the south-west, is a morainic island fringed by beaches of sandy gravel, with a spit with boulders at the western end. A sheltered low embayed coast made up of Quaternary deposits borders Suur and Väike väin, a narrow strait east of Saaremaa Island. The offshore zone is very shallow. There are low marshy areas and occasional shingle beaches on the shores lined with boulders derived from the glacial drift. There are some low scarps cut into till and fluvioglacial deposits. Dunes are either poorly developed or missing at all. Exposure of the coast to wind and wave activities increases south of Kõrgelaid, where there is a recurved spit. Reed beds occupy the sheltered bays, and a moraine ridge has become a spit with boulders at Liu. At Valgeranna, reed has spread onto a sandy beach, a feature that is seen on many Estonian beaches (Bird et al. 1990). A long and wide gently curving beach of fine sand rich in quartz backed by low dunes has formed at the head of Pärnu Bay. Reed beds have spread over parts of the shore, particularly west of the harbour jetties here, too. On the eastern coast of the Pärnu Bay, emergence following the Litorina transgression, which here attained 5–7 m above the present sea level, resulted in the formation of a coastal plain and a shore fringed by boulders, sandy beaches and segments of reed beds. To the south, as exposure to wave activity increases along the coast of Gulf of Livonia,

8.4

beaches are backed by lagoons overgrown with reed. At Heinaste, marshes occupy the shore because the sand drifting north along the coast is being intercepted by a jetty built on the Latvian frontier.

3.  Hiiumaa and Saaremaa The eastern and southern coasts of these two large islands are low and flat, locally marshy, with pebble and cobble beaches, while the northwestern coast of saaremaa has up to 20 m high cliffs cut into Silurian limestone. The nearshore area here is an abrasion platform of bare rock. Pebble and boulder beaches are widespread near the cliffs. Sand and gravel beaches have been derived from eskers and end moraines, and are distributed by longshore drifting (Orviku et al. 1995). Erosion of limestone bedrock has yielded gravel and pebble, which drifts along the shore and accumulates in spits and beach ridges. Some cliff-edged bays are the sides of former large river valleys that descended tens of metres below the present sea level and have been filled with Quaternary sediment, including glacial drift, some of which have been washed up as sandy beaches. Erosion of the cliffs of Silurian limestone at Panga and Ninase has yielded gravel deposits, which have drifted into intervening Küdema Bay to accumulate as bay-head shingle beach ridges. In recent decades, there has been erosion on parts of the shingle beach and prolongation of a spit. Due to increased storminess, the rate of changes has been more rapid in 1975–1990 than in the preceding 15 years (1960–1975). The Harilaid Peninsula in western Saaremaa has changed in outline in recent decades, with the northeastward migration of a protruding spit, Cape Kiipsaare. Historical sequences of maps and air photographs show this migration, and a wooden boat wrecked on the northeast shore about 150 years ago was buried and rolled over by sand and eventually emerged on the southwest shore as the result of erosion during the stormy winter of 1999–2000 (Orviku et al. 2003). The Kiipsaare lighthouse, built in the middle of the spit in 1933, now stands on the southwest shore (>Fig.  8.4.6). Sediment eroded from these shores drifts southward forming beaches, thereby widening the connecting barrier (tombolo) between the peninsula and the rest of Saaremaa. Some of the beaches are backed by low grassy dunes. Erosion during storm surges in 1990 cut back a 4 km long coast of sand and gravel by up to 5 m in southwestern Saaremaa, and this erosion has continued up now. The sand

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⊡⊡ Fig. 8.4.6 Kiipsaare lighthouse that was initially erected in the middle of the cape (about 100 m from the western shore) in 1933 is currently on the shoreline. Erosion has caused it to tilt seaward by 7°.

drifts northeast along the shore from the scarped coast of Tehumardi-Järve and accumulates beside the Nasva harbour breakwaters and in the harbour approach channel during the storms. Some of the finer sediment drifts on into the old waterway Tori harbour (Port of Kuressaare). Locally the shore is strewn with silty sediment. The islands were much smaller during the Ancylus Lake stage and successively enlarged at the Litorina Sea and Limnea Sea stages, with subsequent land uplift of the coastal plain. Each of the earlier coastlines is marked by bluffs and beaches, particularly on Saaremaa and Muhu Islands.

4. The Impact of an Extremely Strong Storm on the Coast of Estonia A cyclone known as Gudrun in the Nordic countries developed above the North Atlantic and traveled over Scandinavia to Estonia on 7–9 January 2005. As a result of high initial levels of the Baltic Sea, the fast-traveling cyclone with a favorable trajectory and strong SW–W winds created a record high storm surge (275 cm) in Pärnu, as well as in many other locations along the west Estonian coast. The streets near the sea in Pärnu were flooded up to the city center. Seven hundred and seventy five houses in an area of 8 km2 were affected and 400 people were evacuated. The January storm induced clearly visible changes in the development of shores and the dynamics of

beach sediment over almost all of Estonia. The precondition for the profound changes observed from this storm was a combination of the absence of protecting ice cover in the sea, relatively high sea level for a long period before the storm, and a very intensive storm surge, taking place over the background of the already elevated sea level. Strong storm waves combined with the high sea level caused substantial changes in the coastal geomorphology of depositional shores. The most exceptional changes occurred in the areas that were well exposed to the storm winds and wave activity—for instance, in Kelba (NW Saaremaa Island), where the high rate of erosion (>3000 m3) resulted in the elongation of a spit by 75 m. The January 2005 storm caused significantly larger changes to the depositional shores in west Estonia than the cumulative effects of ordinary storms over the preceding 10–15-year period (Tõnisson et al. 2008).

References Bird ECF, Martin E, Orviku K (1990) Reed encroachment on Estonian beaches. Proc Est Acad Sci 39:7–12 Martin E, Orviku K (1988) Artificial structures and shoreline of Estonian SSR. In: Walker HJ (ed) artificial structures and shorelines. Kluwer Academic Publications, Boston, pp 53–57 Orviku K, Bird ECF, Schwartz ML (1995) The provenance of beaches on the Estonian islands of Hiiumaa, Saaremaa and Muhu. J Coastal Res 11:96–106

Estonia Orviku K, Granö O (1992) Contemporary coasts. (In Russian, English summary) In: Raukas A, Hyvarinen H Geology of the Gulf of Finland. Estonian Academy Publishers, Tallinn, pp 219–238 Orviku K, Romm G (1992) Litho-morphodynamical processes of Narva Bay. (In Russian, English summary). Proc Esto Acad Sci Geol 41:139–147 Orviku K, Sepp V (1972) Stages of geological development and landscape types of the islets of the West-Estonian Archipelago. Geographical Studies. Academy of Sciences of the Estonian S.S.R., Estonian Geographical Society, Tallinn:15–25 Orviku K, Jaagus J, Kont A, Ratas U, Rivis R (2003) Increasing activity of coastal processes associated with climate change in Estonia. J Coastal Res 19:364–375

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Raukas A (2005) Klint coast. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Netherlands, pp 586–587 Tõnisson H, Orviku K, Jaagus J, Suursaar Ü, Kont A, Rivis R (2008) Coastal damages on Saaremaa Island, Estonia, caused by the extreme storm and flooding on 9 Jan 2005. J Coastal Res 24: 602–614 Vallner L, Sildvee H, Torim A (1988) Recent crustal movements in Estonia. J Geodyn 9:215–223

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8.5 Latvia

Guntis Eberhards . Viktor Brenners

1.  Introduction The coast of Latvia is about 496 km long, extending from the Estonian border southward around the coast of the Gulf of Riga to Cape Kolka (Kolkasrags) on the Kurzeme Peninsula, then southwest past Ventspils, and Liepaja to the Lithuanian border. Much of the Latvian coastal area is an undulating lowland with Pleistocene glacial drift deposits of varying thickness concealing Devonian bedrock (sandstone, dolomite and clay), and Holocene marine sands and gravels, alluvial deposits, aeolian sands, lagoonal peats, clays and gyttja. The coast is generally low-lying, often between 4  and 8 m high. There are extensive sand and gravel

beaches and boulder-strewn shores and nearshore areas, and locally reeds and rushes grow on the shore. Cliffs cut in glacial drift occur along separate sections of the open Baltic coast from Liepaja to Ventspils, bluffs cut in soft marine and aeolian sediments stretch from Ventspils to Kolka, with low cliffs cut in Devonian sandstones only along the eastern coast of the Gulf of Riga near Tuja and Cape Kurmrags. The total length of sandy beaches along the coastline of Latvia is about 240 km, with 80 km classified as having high recreational potential. About 70 km of coastline consists of narrow gravel, shingle and boulder beaches or is overgrown with reeds and rushes. The remainder of the coastline (140–180 km) consists of sandy and pebbly beaches (>Fig. 8.5.1). Sandy beaches are

⊡⊡ Fig. 8.5.1 Types of beaches on the Latvian coast.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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subject to the most intense seasonal and cyclic changes. During individual storm events sandy beaches can be reduced in width by 20–30 m and the volume of sand reduced by 10–30% (Eberhards and Saltupe 1993, 1999). Foredunes have formed where sand has been blown from the beach (>Fig.  8.5.2). Typically they are up to 2.5 m high, but where sand accretion is rapid, as on the southern shore of the Gulf of Riga, they may attain a height of 6 m. Here, along a 56 km stretch of coastline, about 50,000 m3/year of sand is transported by onshore winds from the beach to the foredune. About 125 km of the Latvian coast is eroded during severe north-westerly storms. Within the coastal zone (up to 50 m wide) threatened by erosion there are over 120 structures, including 50 residential buildings, three lighthouses, three cemeteries and minor roads. About 62% of the Latvian coast must be classified as highly ­vulnerable to erosion (Eberhards and Saltupe 1993). In response to ­erosion coastal structures such as rip-rap and  revetments have been built, and presently occupy a  total length of ­about 10 km, or 2% of the Latvian coastline (Eberhards 1998a). There are also artificial

structures of various kinds at and near harbours, towns and fishing villages. Coastal evolution in the eastern Baltic area has been influenced by the Litorina transgression and the ensuing regression (7,000–2,800 years ago). During this transgression there was extensive erosion, producing cliffs and shore platforms and generating large amounts of sand and gravel, with residual boulders where morainic deposits had been dispersed. During the regression sand, gravel and boulders on the emerging sea floor were shaped into beaches, beach ridges and dunes by wave and wind action (Gudelis and Konigsson 1979), forming a modern depositional terrace up to 10 km wide (Straume 1979). Sand also  drifted north from eroding cliffs of glacial drift in Kaliningrad and Lithuania to the beaches of south–west Latvia, eventually forming the depositional foreland at Cape Kolka. Although several rivers drain to the coast between Nida and Kolka their yield of sandy sediment is small. The three largest rivers in Latvia, the Lielupe, Dauvaga and Gauja, have a total sediment yield of up to 13,000 m3/year. Sand is distributed by longshore drifting (>Fig. 8.5.3). ⊡⊡ Fig. 8.5.2 Foredunes and coastal erosion in Latvia.

Latvia  

Postglacial isostatic movements have resulted in uplift of about 5 m in southern Latvia, increasing to 10–12 m in the north of the Kurzeme Peninsula. Behind the sandy beaches there are often foredunes, older parallel foredunes and newer parabolic and active dunes (Eberhards 1998b, 2000), the older ridges to landward having been raised by  tectonic movements. Tilting has continued on either side of a zero isobase that runs SW–NE through Riga (>Fig. 8.5.3), with subsidence of about 1 mm/year in the far south and uplift of a similar amount near the Estonian border. However, these crustal movements have had little effect on recent coastal dynamics. Tides are negligible on the Latvian coast, but sea level is generally higher in winter than in summer. Southerly and southwesterly winds predominate, producing waves from these directions on west-facing sectors of coast, but ⊡⊡ Fig. 8.5.3 The coast of Latvia, showing the pattern and volume (in thousands of cubic metres per year) of longshore drift northward and round Cape Kolka and on either side of the Gulf of Riga, the height of the coastline of the Late Glacial Baltic Ice Lake (numbers in squares) and isobases of land tilting (mm/year). S Saaremaa Island, Estonia, which partly blocks the mouth of the Gulf of Riga.

   8.5

wave energy is generally low, especially in the Gulf of Riga where the nearshore is shallow and boulder-strewn. During storms wave action is stronger, and in storm surges the onshore winds and low barometric pressure may raise sea level by 1–1.5 m, and more than 2 m on the southern coast of the Gulf of Riga. This is when major changes take place along beaches, and large quantities of sand move alongshore (Gudelis 1967; Knaps 1966). Until 1960 a shore ice fringe developed and impeded wave action during winter (December–April), and when it broke up waves drove it onshore, displacing sand and gravel, and even boulders. During the past 40–50 years the winter shore ice l fringe has been reduced. Coastal erosion has accelerated in Latvia in recent decades, partly because of a rising sea level and partly because of increased wind and wave energy resulting from  stronger and more frequent storms (Eberhards and Saltupe 1995). The annual mean sea level rise has decreased during last 30 years, with a total rise of about 8–10 cm (Ulsts 1998). Only in Daugavgriva (at the mouth of River Daugava) has the long-term mean sea level risen uninterruptedly during the last 120 years, increasing by 30 cm. The main reason for this is negative neotectonic movement and land surface settling as result of artesian water extraction.

2.  The Gulf of Riga PARNU

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The Gulf of Riga is a submerged lowland 18,000 sq. km in area with depths of up to 50 m. It was excavated mainly by glacial erosion during the Pleistocene (Gudelis 1970, 1973), and has received Baltic Ice Lake, Ancylus Lake, Litorina and post-Litorina Sea sediments. The present coastline is 239 km long. The east (Vidzeme) coast of the Gulf of Riga from the Estonian border south to Saulkrasti, is about 60 km long. It is an eroded coast, where sandy beaches are interrupted by narrow (up to 20 m) gravelly and bouldery beaches. Nearshore platforms have been cut in Devonian sandstones and clays or in glacial deposits, with occasional boulders and ridges of boulders derived from glacial drift. The cliffs are generally cut in glacial drift sediments, but north of Tuja and near Cape Kurmrags they are up to 4–6 m high, cut in Devonian sandstones and clays, backed by the flat erosional Baltic Ice Lake plain. Along the shore from Ainazi and Salacriva south to the River Svetupe, there are sandy beaches with reeds and rushes and maritime meadows, known locally as “randu plavas”. Longshore drifting of sand is southward at a rate of about 15,000–25,000 m3/year, contributing to the accretion on the south coast of the Gulf of Riga.

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The south coast, from the River Gauja to the River Lielupe, is a sandy depositional coast with beaches 30–50 m wide and foredunes up to 6 m high, behind which are dune ridges up to 20–30 m high, some bearing pine forest while others have drifted inland during recent centuries. Behind the dunes are Litorina Sea lagoonal plains. The dune sand has been partly delivered by longshore drifting from the eastern and western shores of the Gulf of Riga, and partly supplied by the Rivers Daugava and Gauja (Ulsts 1957). There are also many intertidal and subtidal sand bars, some of which overlie gravel ridges. Between the seaside resort of Jurmala and Engure coastal dunes have been eroded, mainly during the severe storm of 1969. The west coast of the Gulf of Riga is a low plain formed during the Litorina Sea phase, mainly depositional, with sandy beaches 20–50 m wide, spits up to 1 km wide and parallel dune ridges, behind which there are lagoonal plains with partly overgrown lakes (Engure, Kanieris, Babite) and wetlands (Eberhards 2000). There is a sandy nearshore zone with up to five sand bars. Locally, in eroded sectors, a shore platform has been cut in Devonian sandstone or in glacial deposits near Engure harbour and between Roja and Kaltene. There are boulders scattered along the shore and in the nearshore zone, narrow gravel and shingle or boulder beaches, in some sectors reeds, rushes and shore meadows near Mersrags harbour and Berzciems village. Sand drifts round Cape Kolka and then southward along the Gulf of Riga shore at the rate of up to 50,000 m3/year. Where breakwaters protrude, as at small harbours such as Engure, Mersrags and Roja, there is some updrift accretion and much downdrift erosion. Up to 2 km of the coastline is affected by downdrift erosion, the erosional scarp retreating at up to 150 m/year (Eberhards 1998a). Erosion was severe during a hurricane in 2005 (Eberhards and Saltupe 2006). The sandy coast curves out at the northern end to Cape Kolka.

3.  The Coast of southwest Latvia Cape Kolka is a large sandy foreland with about 200 parallel dune ridges, formed as a result of accretion of sand delivered by longshore drift from the south along the coastline established at the end of the Litorina Sea regression (>Fig.  8.5.4). At the northern end there is an extensive nearshore shoal in front of the Kolka lighthouse. It has been estimated that the volume of sand drifting along the shore north from Ventspils to this point is about a 0.75–1 million m3/year (Gudelis and Jemeljanov 1976; Knaps 1966). Irbe Strait separates Cape Kolka from Saaremaa Island in Estonia. The coast west of Cape Kolka is a typical sandy

⊡⊡ Fig. 8.5.4 Geomorphology of Cape Kolka. 1 – parallel dune ridges, 2 – blowouts, 3 – cliff-top dunes, 4 – complex parabolic dunes, 5 – heights in metres.

depositional plain with parallel dune ridges and foredunes up to 6 m high, fringed by sandy beaches up to 60 m wide and a sandy nearshore zone with up to five sand bars. Sand drifting from the south has accumulated mainly in the nearshore zone of Irbe Strait, with a relatively small amount arriving on the beach and foredunes (Ulsts 1998). The 257 km west-facing coast south from Liepaja has sand and gravel beaches 30–110 m wide with scattered boulders. The beaches are backed by dunes up to 5 m high (>Fig. 8.5.5), and in some sectors there are lagoons and swamps behind the dune fringe. One of the lagoons has an outlet close to Liepaja. The coastline has been straightened by strong erosion during the past 50 years, with cliffs retreating up to 200 m (Eberhards 1998a). At Cape Ovisi the coastline turns sharply to the southwest. During the last century breakwaters at Ventspils harbour, where the River Venta flows into the sea, have interrupted the longshore sediment drift from the south and there is a broad accretion (300–700 m in width)

Latvia  

   8.5

⊡⊡ Fig. 8.5.5 Sparsely vegetated dunes behind the beach near Mazirbe.

⊡⊡ Fig. 8.5.6 The coast near Pavilosta, showing a steep slope cut into glacial drift deposits behind a beach of sand and gravel with ­occasional boulders derived from the glacial drift.

with foredune ridges on the southern side. South of Ventspils to Jurkalne lighthouse and Pavilosta harbour is a 30 km sector where cliffs up to 20 m high have been cut into glacial, interglacial marine and Baltic Ice Lake sediments. These cliffs are retreating by up to 2.5 m/year, and during severe storms thisfigure can reach 10–15 m. Near the village of Jurkalne the cliffs

have retreated 100–150 m during the past 60 years. Beaches are usually narrow, up to 15–30 m in width, with a thick cover of coarse sand and shingle. The nearshore slope is cut steeply into glacial deposits, with boulders and one or two sand bars. An erosional coast dominates south from Pavilosta to Liepaja, where cliffs up to 10 m high have been cut into

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⊡⊡ Fig. 8.5.7 Cliffed dune along the backshore of a recently eroded beach in southern Latvia. (Courtesy Liepajas udens.)

glacial drift, fringed by narrow sandy and gravelly beaches (>Fig.  8.5.6). Erosion has accelerated here because sand ­drifting northward has been intercepted by breakwaters at the port of Liepaja seaport. There is accretion on the southern side forming a beach of fine sand with foredunes, and a broad nearshore shoal with sand bars. South of Liepaja the coast has fine sandy beaches 30–80 m wide with foredune ridges, prograded by accretion of sand that has drifted alongshore from the south. The beaches show alternations of cut and fill, with erosion during stormy periods leaving backshore dunes cliffed (>Fig. 8.5.7) and accretion in calmer weather. Behind the dunes are lagoons and marshy plains. The largest of these lagoons are Lake Liepaja and Lake Tosmare with an outlet close to Liepaja, and Lake Pape, with an outlet close to Pape. Towards the Lithuanian border are narrow (15–30 m) stretches of gravel and shingle beach, extending for 15 km. Behind the beaches is a major complex dune ridge up to 30–34 m high, formed during the Litorina Sea stage. Near Cape Bernati and Cape Mietrags there has been increased erosion of the beach and dune coast during the last decades.

References Eberhards G, Saltupe B (1993) Sea coast monitoring of Latvia. Envi­ ronment monitoring of Latvia, 3:46

Eberhards G (1998a) Harbours and coastal processes in Latvia. Abstracts of papers and posters. INQUA Commission of glaciations. Field symposium on glacial processes and quaternary environment in Latvia. University of Latvia, Riga, pp 15–17 Eberhards G (1998b) Coastal dunes in Latvia. Environmental perspectives of Southeast Baltic coastal areas through time, Field Guide, Riga, pp 18–25 Eberhards G (2000) Litorina sea coastal formations and the origins of stone age habitation on the shore of the Gulf of Riga in northern Kurzeme: geological background. Archaeol Ethnogr 10:211–222 Eberhards G, Saltupe B (1995) Accelerated coastal erosion – implications for Latvia. Baltica 9:16–28 Eberhards G, Saltupe B (1999) The sea coast processes monitoring in Latvia – experiment and practice. Folia Geographica 7:1–10 Eberhards G, Saltupe B (2006) Coastal erosion in the Gulf of Riga caused by hurricane Erwin in 2005. Abstracts of the ninth marine geological conference (The Baltic Sea Geology), Latvia, 19–21 Gudelis V (1967) Morphogenetic types of the Baltic Sea coasts. Baltica 3:123–145 Gudelis V (1970) Main features of geology and bottom topography of the Mid-Baltic Sea. Baltica 4:103–113 Gudelis V (1973) Relief and quaternary of the east Baltic region. Mintis Publishing House, Vilnius Gudelis V, Jemeljanov J (1976) Geology of the Baltic sea. Mokslas Publishers, Vilnius Gudelis V, Konigsson LK (eds) (1979) The quaternary history of the Baltic. Almquist and Wiksell, Uppsala, Sweden Knaps R (1966) Eastern Baltic longshore sediment drift. Evolution of sea coasts in fluctuating conditions of tectonic movements. Tallin 21–39 Straume J (1979) Geomorphology. Geological structure and minerals in Latvia. Riga, p 427 Ulsts V (1957) Morphology and development of marine accumulation area in the Gulf of Riga head. Academy of Sciences of the Latvian SSR, Riga Ulsts V (1998) Latvian coastal zone of the Baltic Sea. Riga, p 96

8.6  Lithuania

Eric Bird

1.  Introduction The coast of Lithuania is about 100 km long. The Kaliningrad border to the south bisects the inner shore of the Kursiu Marios (Kurisches Haff), a large freshwater lagoon separated from the Baltic Sea by the long barrier spit that extends from the Kaliningrad coast to the south. The lagoon has an outlet at Klaipeda, at the northern end, and beyond this the coast north to the Latvian borderis generally lowlying, with sand and gravel beaches bac­ked by dunes and some low bluffs cut in glacial drift (> Fig. 8.6.1). The eastern Baltic area is an extension of the East European plain, and has a mantle of Pleistocene glacial

⊡⊡ Fig. 8.6.1 The Lithuanian coast. (Courtesy Geostudies.)

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drift deposits over Jurassic, Cretaceous and Lower Tertiary bedrock. There are morainic deposits and outwash sands, with a partial covering of Holocene alluvial and lagoonal deposits, with peat and wind-blown sand. The Lithuanian coastline was shaped after the Litorina Sea transgression and the following regression (7,000– 5,000 years ago). During the transgression there was rapid erosion of cliffs cut in glacial drift, producing large amounts of sand and gravel. During the regression, the emerging coastline straightened and spits, deltas and coastal dunes were formed. The postglacial period was characterised by glacioisostatic and tectonic crustal movements, which raised the land surface north of the Nemunas (Njemen) delta and resulted in subsidence in Kaliningrad to the south. The emerged Holocene beach ridges seen in Estonia and Latvia are not present in Lithuania. In recent decades there has been slow subsidence along the Lithuanian coast, increasing southward to about 1 mm/year at the Kaliningrad border. These crustal movements have had limited influence on recent coastal dynamics, their effects being masked by longshore sediment drift and wind action (Gudelis and Kønigsson 1979). The tide range is negligible on the Lithuanian coast, but sea level is generally higher in winter than in summer. Westerly and southwesterly winds predominate, producing waves from these directions on west-facing sectors of coast, but wave energy is generally low, especially where the nearshore is shallow and boulder-strewn. Westerly winds generate larger waves across the longer westward fetch. During storms wave action is stronger, and in storm surges the onshore winds and low barometric pressure may raise sea level more than 2 m. This is when major changes take place along beaches, and large quantities of sand move alongshore, mainly from south to north. During the winter (December–April) a shore ice fringe develops and impedes wave action, but when it breaks up waves may drive it onshore, displacing sand and gravel, and even boulders. The sand and gravel beaches along the Lithuanian coast consist mainly of sediment derived from cliff erosion on the Sambian Peninsula, a coastal plateau, rising 40–60 m above sea level in Kaliningrad to the south.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.6.2 The sandy coast of the Baltic Sea along the Kursiu Nerija barrier spit. (Courtesy Geostudies.)

⊡⊡ Fig. 8.6.3 The inner shore of the Kursiu Nerija spit showing dunes spilling towards the lagoon. (Courtesy Geostudies.)

The cliffs are cut into Tertiary sands and clays and a thin overlying cover of Quaternary morainic deposits. Sand thus derived has drifted northward to build the barrier spit fronting the Kursiu Marios lagoon, and on beyond Klaipeda to the Latvian border and beyond. The volume of drifting sand and gravel is large, of the order of half a million cubic metres per year. During the second half of the twentieth century coast protection works were introduced to sectors of cliffs on the Sambian Peninsula, and the supply of sand drifting northward began to decline. As a consequence, coastal erosion has accelerated in Lithuania in recent decades, a change that may be partly due to the slight sea level rise accompanying land subsidence.

2.  The Lithuanian Coast From the Latvian border south to Klaipeda the coast is bordered by sand and gravel beaches 30–110 m wide, backed by foredunes. In some places behind the foredunes there are relics of small coastal lakes and swamps. At intervals along the shore are boulders derived from glacial moraine deposits, the finer sand and gravel having been washed away by wave action. Beyond Klaipeda is the mouth of the Kursiu Marios lagoon, the narrow Strait of Klaipeda separating the Kursiu Nerija spit from the mainland at its northern end. This barrier spit, the longest in the Baltic (98 km long, up to 4 km wide), stretches in a gentle curve (> Fig. 8.6.2) south

Lithuania

to the cliffed coast of the Sambian Peninsula. It is bordered seaward by a sandy beach up to 80 m wide, backed by grassy foredunes. There are also active dunes that are locally more than 60 m high, some of which are spilling over into the lagoon (> Fig. 8.6.3). The inner shore of the spit is irregular, with a series of cuspate and lobate forelands. Behind the Kursiu Marios lagoon the inner shore is low  with reed and rush swamps and alluvial deposits,

8. 6

especially on the delta of the Nemunas River, just north of the Kaliningrad border.

Reference Gudelis V, Kønigsson LK (eds) (1979) The quaternary history of the Baltic. Almquist and Wiksell, Uppsala, Stockholm

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8.7  Kaliningrad

Alexey Porotov

1.  Introduction The Kaliningrad coastline (>Fig. 8.7.1) in the Kaliningrad administrative region or oblast on the southeastern Baltic Sea extends for 220 km. It consists largely of Pleistocene glacial drift deposits overlying Tertiary formations, which outcrop in the cliffs of the Sambian Peninsula in central Kaliningrad, and Holocene sediment including beaches and dunes along the coast. To the north of the Sambian Peninsula is the Curonian barrier spit, 98 km long and up to 4 km wide, which extends into Lithuania, where it is known as the Kursiu Nerija spit. On the southern side is the similar Baltiisk barrier spit, over 60 km long, extending northward from Poland, where it is known as the Vistula (Wislany) spit. These two long barrier spits have sandy beaches up to 80 m wide, backed by grassy foredunes and active dunes that are locally more than 60 m high, some of which are spilling over into the backing brackish lagoons, the Curonian Gulf (Kursiu Marios) to the north and the Kaliningrad Gulf (Vistula Lagoon) to the south. The spits have received sand from the eroding cliffs of the Sambian Peninsula and from the sea floor, but the relative importance of these two sources has not been determined. The coastline of the SE Baltic was shaped after the Litorina Sea transgression and the following regression (7,000–5,000 years ago). During the transgression there was rapid erosion of cliffs cut in glacial drift, producing large amounts of sand and gravel. During the regression, the emerging coastline straightened as barrier spits and coastal dunes were formed, enclosing coastal lagoons. The postglacial period was characterised by glacioisostatic and tectonic crustal movements, which resulted in uplift in Lithuania and Latvia and slow subsidence in Kaliningrad, the zero isobase running through the Nemunas delta at the border between the Kaliningrad region and Lithuania. In recent decades the Kaliningrad coastline has been subsiding at about 1 mm/year, but the effects have been masked by longshore sediment drift and wind deposition (Gudelis and Kønigsson 1979). Tides are negligible on the Kaliningrad coast, but sea level is generally higher in winter than in summer. Westerly

and SW winds predominate, producing waves from these directions on the Sambian Peninsula and the Curonian barrier spit, but wave energy is generally low, especially where the nearshore is shallow and boulder-strewn. Westerly winds generate larger waves across the longer westward fetch. During storms wave action is stronger, and in storm surges the onshore winds and low barometric pressure may raise sea level more than 2 m. This is when major changes take place along the beaches. During the winter (December–April) a shore ice fringe develops and impedes wave action, but when it breaks up waves may drive ice floes onshore, displacing sand and gravel, and even boulders. The sand and gravel beaches along the Kaliningrad coast consist mainly of sediment derived from cliff erosion on the Sambian Peninsula, and carried in from the floor of the Baltic Sea. During the second half of the twentieth century coast protection works were introduced to sectors of cliff on the Sambian Peninsula, and the supply of drifting sand began to decline. As a consequence, beach erosion has accelerated in recent decades, a change that may be partly due also to a slight sea level rise accompanying land subsidence.

2.  The Kaliningrad Coastline The outer (Baltic) shore of the Kursiu barrier spit has some accreting sectors, notably near the Lithuanian border, where the beach is wide and onshore winds are delivering sand to backshore dunes, but to the south there are eroding sectors with cliffs cut into the dunes (>Fig. 8.7.2), in places exposing the underlying glacial drift (>Fig.  8.7.3). There are narrow sectors that have been overwashed by storm surges. The inner shore of the spits are irregular, with a series of cuspate and lobate forelands, and the two freshwater lagoons, Curonian Gulf to the north and Kaliningrad Gulf to the south, have landward shores that are low-lying, with reed and rush swamps and alluvial deposits. The hinterland is an undulating lowland of glacial deposition with morainic hills and marshy hollows. Pine forests are extensive on dunes and sandy glacial drift.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.7

Kaliningrad

⊡⊡Fig. 8.7.1 The Kaliningrad coastline. (Courtesy L. A. Zhindarev.)

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⊡⊡Fig. 8.7.3 Scarp cut in Pleistocene sediment in the southern part of the Curonian Spit. In the lower part of the scarp glacial morainic loams are exposed, overlain by an old soil layer, capped by dune sand. (Courtesy L. A. Zhindarev.)

⊡⊡Fig. 8.7.4 The northern coast of the Sambian Peninsula, showing cliffs cut in Pleistocene glacial and fluvioglacial sediment behind eroding beaches that are partially retained by anti-landslide structures and wooden groynes over 100 m long. (Courtesy L. A. Zhindarev.)

The Sambian Peninsula is a coastal plateau, rising 40–60 m above sea level. It is bordered by cliffs are cut in Tertiary sands and clays overlain by a thin cover of Quaternary morainic deposits. They contain deposits of amber (fossilised resin) which have been exploited over many centuries. The north-facing cliffed coastline is slightly sinuous, with narrow beaches of sand and gravel and some

cobbles and boulders. The cliffs increase in height westward, and show slumping, dissection by ravines and erosion by groundwater seepage, all of which deliver sand and gravel to the shore (>Fig. 8.7.4). The nearshore area is shallow, strewn with boulders and gravel. At Brusterort Point the Sambian Peninsula coastline turns southward, and the west-facing cliffs diminish in height (>Fig. 8.7.5).

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⊡⊡Fig. 8.7.5 The western coast of the Sambian Peninsula with a cliff up to 7 m high, fronted by waste from a nearby amber-producing plant. (Courtesy L. A. Zhindarev.)

Brusterort Point marks a divergence of longshore drifting. Sand derived from the cliffs has drifted northward to build the barrier spit fronting the Curonian Gulf, and on beyond the Latvian border, and also southward along the shore of Gdansk Bay. The volume of northward drifting sand and gravel is large, of the order of half a million cubic metres per year, while that to the south is somewhat less. As the cliffs fade out south of Brusterort Point a sandy beach widens on the outer shore of a spit that extends to Baltiysk (formerly Pilau). Outcrops of amber-bearing earth are mined at Yantarny. The seaside resorts of Svetlogorsk and Zelenogradsk stand on lobate promontories fringed by sandy beaches. Baltiysk is a naval port on the northern side of the entrance to the Kaliningrad (Vistula) Lagoon.

It is also the outport of Kaliningrad, which was formerly Kønigsberg, an historic seaport at the mouth of the Pregolya River, linked by a canal dredged across the shallow Vistula Lagoon. The barrier spit to the south resembles the Kursiu barrier spit, and the inner shores have sectors of reed fringe. Balga is a fourteenth century fort at North Cape, on the eastern shore of Kaliningrad Lagoon, 40 km south of Kaliningrad.

Reference Gudelis V, Kønigsson LK (eds) (1979) The quaternary history of the Baltic. Almquist and Wiksell, Uppsala

8.8  Poland

Karol Rotnicki · Joanna Rotnicka

1.  Introduction The Baltic coastline of Poland, excluding the Puck Bay coast along the Hel Spit and the inner coastline of Vistula Haff in the east and Szczecin Haff in the west, is 469 km long. It may be divided into three sections: eastern, middle and western. The eastern coast, 172 km long, extends from the Russian border on the Vistula Barrier in the east to the northernmost point of Poland, Cape Rozewie (54° 50' N) in the west. It runs round Gdansk Bay and Puck Bay, and contains both sides of the Hel Spit. Cliffs constitute 18.3% of this coast and are particularly prominent between Wladyslawowo and Jastrzebia Gora. The middle coast, 179 km long, runs westward and ends some kilometres beyond the holiday resort Mielno. It contains six coastal lakes and three lakes of glacial origin: Zarnowiec (a subglacial tunnel valley), Gardno and Wicko (both shallow basins in the former ground moraine). The western coast, 118 km long, extends westward as far as the German border, some kilometres beyond the port and holiday resort of Swinoujscie. It is the straightest section of the Polish coast, and 42% of it is cliffed. In the western part there are two islands, Wolin and Uznam (partly German), which enclose the estuarine lagoon basin of Szczecin Haff to the south. The Polish coast lies in the temperate zone. The mean annual temperature on the Polish coast is 7.1°C at Cape Rozewie and 8°C in Swinoujscie. Mean temperatures in January ranges between −0.6°C on the western coast (Swinoujscie) and −2.4°C on the eastern coast (Swibno, Vistula Spit). July is the warmest month (16.6°C–17.0°C) on the western coast and August (16.1°C–17.0°C) on the middle coast, east of Leba. The highest average annual rainfall is recorded on the middle coast between Darlowo and Leba (650–700 mm), and is less (550–570 mm) on other parts of the coast. The coast is dominated by southwesterly and westerly winds, which are strongest in autumn and winter and weaker between May and August. Most days with strong winds and storms occur on the middle and eastern coast (Ustka, Leba, Rozewie, Hel, Gdansk) in the period between November and January.

The most frequent wind-generated waves in the southern Baltic are less than 1 m high. The highest waves reach 7–8 m, but waves higher than 3 m occur very seldom. Storm waves in the offshore zone reach 1 m above mean sea level, and as they cross steeply sloping nearshore areas they rise to 2 m. Tidal amplitude on the Polish coast is very low, 3.6 cm in Kolobrzeg and 3.3 cm in Wladyslawowo. The coast may be thus classified as wave-dominated since tides do not have much influence on the morphological evolution of the coast. Prevailing westerly winds generate an eastward longshore current that is pronounced along much of the Polish coast (> Fig. 8.8.1). Sediment (mainly sand) in the nearshore zone drift eastward (Mojski 1995a), except on the coast between Koszalin Bay (middle coast) and Swina Gate (western coast), where the drifting is westward, and in the area of Puck Bay and Puck Lagoon, where longshore currents and sand transport are governed by local coastline configuration. Much of the Polish coast is developed in Quaternary deposits, with an underlying surface of Cretaceous marlstones and siltstones in the deeper part of the Gdansk Bay, Tertiary quartzose sands, silts, and clays occurring in the area between the meridian crossing Gdansk and the western end of the middle coast in the vicinity of Mielno, and Jurassic limestones, marlstones, sandstones, and Cre­ taceous marlstones in the western segment (Mojski 1995a). Of these only the Tertiary deposits outcrop on the coast above sea level, exposed in the lower parts of cliffs close to Gdynia and Wladyslawo. At the latter site the Tertiary and Qua­ternary deposits have been glacitectonically deformed.

2.  Quaternary Evolution The Quaternary deposits offshore comprise glacial tills of the Odra (Saalian) and Vistula (Weichselian) Glaciations, glaciofluvial and fluvial sands and gravels, and glacial lake silts and varved clays. Similar sediment are exposed in cliffs, where in several places they have been glacitectonically disturbed, as at Debina on the middle coast and Wolin

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.8.1 Geology, landforms and processes on the Polish coast. 1 – axis of the Tornqvist-Teisseyre zone, 2 – predominant direction of the longshore drift, 3 – direction of silt and clay transport from cliffs, 4 – mean annual rate of cliff erosion (metres per year in numerator) and cliff height (metres in denominator), 5 – direction of main Pleistocene ice sheet flow, 6 – rate of sea level rise on the Polish Coast 1951–1990. (Courtesy Geostudies.)

Island between Miedzyzdroje and Dziwnow on the western coast. On the middle coast, where lagoons and barriers have well developed, the thickness of the Holocene marine sands and silts along the coast is about 2–5 m. In the zone of present barriers and coastal lakes these marine Holocene deposits, together with lagoon silts, are 8–15 m thick. Amber deposits are found in muddy sediment beneath beaches and dunes near Gdansk, and are harvested by trawling nets. During the recession of the last ice sheet the presentday shallow part of the southern Baltic was land dominated by newly formed glacial relief, characterised by chains of end moraines, ground moraine plains, outwash plains and ice marginal valleys. At the beginning of the Holocene the transgres­­ sion  of  the southern Baltic Sea began. During Early Holocene time the rate of relative sea level rise was high, ranging from 32 mm/year to 34.5 mm/year. It resulted mainly from the glacio-isostatic uplift of Scandinavia, ­centred on the Bothnian Gulf. Eventually the role of glacio-isostatic uplift decreased, and in the Atlantic Period the eustatic factor may have become pronounced. Because

of spatial variation in glacio-isostatic movements the maximum extent of the southern Baltic on the present coast was recorded at different times in different places (Uscinowicz 2003). In the area of Vistula Haff, lagoon deposits began to form about 6,000 years ago (Uscinowicz 2003), but in the Puck Lagoon such deposits appeared about 5,500 years ago. On the middle coast, in the northern part of Gardno-Leba Lowland, marine and lagoon deposits appeared for the first time between 7,800 and 7,500 years ago, while on the western coast in Szczecin Haff evidence of marine influence is younger than 6,850 years. The ­glacio-isostatic factor has also caused the southernmost former coastline of the Baltic to lie at various heights, 0.5 m above sea level in Kluki village in the Gardno-Leba Lowland (middle coast) and 10 m below sea level in Gdansk Bay.

3.  Geomorphology The configuration of the southern Baltic coastline has been shaped by the Holocene marine transgression

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advancing over a glaciated landscape with Pleistocene deposits. The advancing sea submerged depressions of former ice marginal valleys and river valleys as well as morainic plateaux, and end moraines. Submergence of valley mouths produced an embayed coast, while cliffs were cut into elevated morainic plateaux and end moraines. Longshore drifting of sandy material from eroded cliffs towards embayments built up barriers and spits, enclosing bays as lagoons, as in the Puck Bay (eastern coast) and in the area of Swina Barrier (western coast). The major features of the Polish coast are related to geological and glacial factors. The eastern and middle sections of the coast are located within the Pre-Cambrian crystalline Eastern European Platform, whereas the western coast is within the Palaeozoic Western European Platform, the two being separated by the TornqvistTeisseyre zone. There are regional and local structural blocks related to a system of joints and faults, and differential uplift and subsidence of these blocks formed several elevations and depressions in the basement. Recently such vertical movements of the earth’s crust are still observed on the southern Baltic coast. The middle coast, situated within the so-called Leba elevation, is rising between 0.1 and 0.3 mm/year, while further south there is subsidence of up to 2.5 mm/year in the east and 0.5–1.0 mm/year in the west. The outline of the Polish coast formed over the past  15,000 years – during the Last Scandinavian ice sheet retreat and the Holocene period. These features have resulted from an interaction of glacial deposition, the effects of glacial meltwater and extraglacial rivers in  marginal zones of the retreating ice sheet, glacio-­ isostatic recovery of the earth’s crust, eustatic sea level changes, and climatic factors such as wind and wave regime, storm surges, and the direction of longshore sand transport. Along the coast of Poland several types of coast may be distinguished (Rotnicki and Borowka 1991). Cliffs and bluffs on the Polish coast are cut in ground morainic plateaux and end moraine hills. Cliffs cut in morainic plateaux are 10–20 m high in glacial till of the last (Weichselian) glaciation (> Fig. 8.8.2). The till is often bipartite and composed of ablation sandy till in the upper part and lodgement till in the lower part (Dobracka and Dobracki 1995). Intra-morainic glaciofluvial sands, silts and glacial lake clays are also often exposed in the lower parts of cliffs. Locally Pleistocene sediment are capped with some metres of Holocene aeolian sands, which contain 2 or 3 fossil soil horizons. The highest cliffs reach 25–80 m, and are cut in glaciotectonically thrusted end moraines. Deformation took

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⊡⊡ Fig. 8.8.2 Cliff and shore platform facing Gdansk Bay at GdyniaOrlowo. The cliff has been cut into glacial till, with beds inclined by glaci-tectonic deformation. It is fronted by a cobble and boulder beach, with a variety of erratics derived from Scandinavian regions. The rate of cliff retreat attains 0.8–1.0 m/year. (Courtesy J. Michniewicz.)

place during the youngest (so-called Gardno) phase of the Weichselian, 14,500–14,700 years ago. Cliffs 30–50 m high occur on the eastern coast between Wladyslawowo and Jastrzebia Gora, 25–35 m high on the middle coast around Debina and 50–80 m high on Wolin Island on the western coast (> Fig. 8.8.3). These cliffs are cut in glacial till, and at their base wave action has produced an erosion platform covered with pebbles and boulders. Cliff retreat ranges from 0.4–2.3 m/year. The most rapid recession is on the cliffs near Gdynia (0.8–1.0 m/year), Debina (0.2–2.3 m/year), Jaroslawiec (1.4–2.1 m/year), Trzesacz (0.2–1.1 m/year), and on Wolin Island (0.9 m/ year) (Subotowicz 1995). The erosion rate is probably slower than in the past, when the coast was not protected. Historical data show that the average rate of cliff retreat at Trzesacz

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over the past 500 years amounts to 3.0 m/year (Dobracka and Dobracki 1995). Barriers front lagoons on 43.6% of the middle coast of  Poland. They are not present on the eastern coast, but  occupy 8.7% of the western coast. On the middle coast  there are six coastal lagoons (Lakes Sarbsko, Leba, Dolgie, Kopan, Bukowo and Jamno) situated in former embayments, which have been cut off from the sea by sandy barriers (> Figs. 8.8.4 and > 8.8.5). These embayments were gradually transformed into lagoons connected with the sea

to a varying degree, and eventually into fresh-water lakes. Barrier stabilisation took place in the Younger Holocene, over the past 3,500 years (Borowka and Rotnicki 1995). In some cases newly created coastal lakes were later divided into smaller ones by washover fans formed during episodes of extremely strong storm surges. Large washover fans formed in this way protruded into a large lake that formerly existed near Leba village, cutting it into three: Lakes Sarbsko, Leba and Dolgie. Barriers are built of marine sands, 6–10 m thick, which are underlain by lagoon-lacustrine sediment. ⊡⊡ Fig. 8.8.3 Bluffs and cliffs on Wolin Island cut in push moraine of the Gardno Phase. The cliffs are cut back every few years. (Courtesy W. Stepien.)

⊡⊡ Fig. 8.8.4 Leba Barrier and Leba Lake in the Slowinski National Park. In the central part there is a field of active dunes, and to east and west are areas of temporarily fixed dunes. Formerly the dunes were covered with oak forest (3,000–2,000 bp), then beech forest (2,000–1,000 bp), which since 1,500 bp has been gradually replaced by pines. (Courtesy W. Stepien.)

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⊡⊡ Fig. 8.8.5 Bukowo Lake in the foreground and Jamno Lake in the background are former lagoons cut off from the sea by a narrow barrier. The barrier is wider (up to 200 m) where washover lobes protrude into the lake and are quickly overgrown with reeds. The lakes are separated by a small river delta and washover fans. (Courtesy W. Stepien.)

⊡⊡ Fig. 8.8.6 Cross-section of foredunes and mobile dunes on the Leba Barrier and the Leba Barrier between Cape Czolpino and Leba, showing the sandy beach and a foredune which has been cliffed during a storm surge, then modified by aeolian processes.

There is evidence that they were transgressive, shifting landward during the Holocene marine transgression. Along the Polish coast two types of barriers (narrow and wide) can be distinguished. (Rotnicki and Borowka 1991) Narrow barriers, less than 0.5 km wide, capped

by  a  single foredune and fronted by a rather narrow beach  (20–30 m wide) are found be­­tween Mielno and Jaroslawiec, cutting off Lakes Bukowo, Jamno (> Fig. 8.8.5), and Kopan from the sea. They are being ­occasionally eroded by storm surges. Wide barriers ­(1.0–2.5 km

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⊡⊡ Fig. 8.8.7 The western part the mobile dune field at Cape Czolpino. The forest and small lake are slowly being buried by the migrating dunes.

⊡⊡ Fig. 8.8.8 Peat 3,000–2,000 years old now outcropping on the beach accumulated in inter-dune hollows when the barrier coastline was situated some kilometres further north.

wide) are capped by a complex of parabolic and ­barchan-like dunes. The dunes are generally 20–30 m high, the highest ones reaching 40–56 m on the Leba Barrier (> Fig. 8.8.4), which is 48 km long and extends between Osetnica in the east and Rowy in the west (middle coast). There are up to four foredunes, 4–15 m high, behind the beach (> Fig. 8.8.6). The youngest foredune is covered with a pioneer grass Ammophila arenaria, the older with mountain dwarf pine (Pinus mughus) and the oldest with Pinus silvestris. To the south (i.e. landward) of the foredunes are 7 sq km of arcuate, parabolic and barchan-like dunes (> Fig. 8.8.7). On the Leba Barrier they are partly forested, but west of Leba village there are nine barchan-like mobile dunes, rising to 42 m above sea level. Generally the dunes migrate east and southeast at an average of 10 m/year (Borowka 1990). Within the dunes several fossil soils of different ages (up to 3,000 years bp) may be traced, indicating phases of at least local stability with a vegetation cover, and allowing reconstruction of the history of interplay between aeolian processes and vegetation. Dune formation began about 3,000–3,300 years ago. Locally on the beaches bordering barriers there are outcrops of peat (> Fig. 8.8.8), lagoon silt with a marine malacofauna and the remains of fossil oak forest with trunks in growth position, indicating coastline recession and barrier migration (Rotnicki and Borowka 1995). Some of the lagoons have developed into shallow lakes with low swampy and sandy banks colonised by Scirpus and Phragmites. The total area of these lakes is 196 sq. km,

Poland

with the largest, Leba Lake, 71.4 sq. km. Generally their present depth ranges between 1.5 and 3 m and in a few places reaches 6 m. On the middle coast there are also some ground moraine lakes (Lakes Gardno, Wicko, Resko and Liwia) that were not former lagoon lakes. They formed during the decay of the last ice sheet and at the beginning of the Holocene they were far inland from the southern Baltic coastline, but with the southward migration of this coastline they are now close to the coast (> Fig. 8.8.9). Spits are best developed on the eastern coast, where they constitute 65% of its length. They have formed as the result of eastward longshore drifting. The largest is the Vistula Spit, 96 km long and 0.5–3.0 km wide, extending from east of the Dead Vistula mouth, one of four distributaries within the Vistula delta (> Fig. 8.8.10), and running as an arc as far as the Sambian Peninsula. Its northeastern sector, 18 km in length, belongs to Kaliningrad. Originally the Vistula Haff extended further south and west, but growth of the Vistula delta filled the western part of the haff with deltaic sediment. The Vistula Spit is built of marine sand up to 10–12 m thick, and has three dune generations: (a) inner 4 m high, longitudinal brown dunes located mainly in southern part of the spit, (b) up to 25 m high longitudinal and parabolic yellow dunes in the middle part and (c) up to 4 m high white foredunes on the north coast (Mojski 1995b). The more famous Hel Spit is about 34 km long (> Fig. 8.8.11). It extends east from the Swarzewo morainic

⊡⊡ Fig. 8.8.9 Wicko Lake, a morainic lake situated close to the coastline. The narrow isthmus looks like a barrier but is built of the same deposits as the surrounding morainic plateau. When the lake formed during decay of the last ice sheet at the beginning of the Holocene it was far inland. (Courtesy W. Stepien.)

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plateau (>Fig. 8.8.12) and separates Puck Lagoon and the Puck Bay from the sea. Its 15 km western section is relatively flat and extremely narrow (only 100–300 m), and its single foredune has been completely eroded in few places by increasingly frequent storm surges during the past 20 years. In order to protect the spit the foredune and the beach have been artificially restored by periodic nourishment, using sand dredged from the Puck Bay. East of the village of Kuznica the spit gradually broadens to 3 km at the eastern tip. The relief of this part is generally more differentiate, with foredunes, parallel to the beach, up to 20 m in height and backed by several oblique dune ridges. Unlike the western part of the spit, the eastern sector is becoming wider and longer by sand accretion. Originally the Hel Spit was a barrier chain, composed of several small islands (Tomczak 1995), but in the past 300 years storm surges have narrowed its western part, and the southern coast of the spit has become irregular, with washover fans. The third and shortest spit has recently been built up along the border between Puck Bay and the Puck Lagoon (>Fig. 8.8.12). It begins as the Rewa Spit in the southern part of the bay and continues toward the north as the Seagull Shallow. Seasonally, during the low water table in the Puck Bay, parts of this shallow area emerges as a chain of small and flat islets extending as far as the Hel Spit. This landform has resulted from complicated cell-like water circulation in Puck Bay and the Puck Lagoon. A spit complex with a total length of 18 km occurs around Swina Gate on the western coast (> Fig. 8.8.13).

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⊡⊡ Fig. 8.8.10 The northern part of Vistula Delta, close to the small town of Swibno. The forested area is on the Vistula Spit, which has been broken in three places by former and present Vistula mouths. An artificial channel was cut in 1895. (Courtesy W. Stepien.)

⊡⊡ Fig. 8.8.11 The western part of Hel Spit from Puck Lagoon to Cape Rozewie, the northernmost point of Poland (54° 50' N). The irregular inner coastline of Hel Spit has been formed by sandy washover lobes deposited by storm surges. The dark hollow in the lagoon indicates an area from which sand has been pumped across the spit to renourish the beach. (Courtesy W. Stepien.)

It started to develop in the Younger Holocene, about 5,000 years ago, when the marine transgression reached its ­maximum (Musielak and Osadczuk 1995). The pattern of foredunes shows the evolution of the spit, which has formed at the convergence of long­­shore drifting: westbound nourished by sandy sediment de­­rived from the Dziwnow-Kolobrzeg cliffs, and eastbound ­supplied with sediment from cliffs on the eastern German coast

(Musielak and Osadczuk 1995). The developing spit ­complex gradually cut off the Odra estuary from the Pomeranian Bay forming the Szczecin Haff, a basin which has both estuarine and lagoonal features. The Odra flows to the sea across the Szczecin Haff and through a short channel called Swina River, but during storm surges inflow through this channel forms a delta prograding back into the Szczecin Haff.

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About 130 km of the Polish coast borders alluvial l­owlands and undulating glacial outwash plains as well as former ice marginal and river valleys. The plains are often separated from the beach by one or two foredune ridges, which in many places have been cliffed. In certain areas this beach and foredune system is attached to dunes that formed far inland, before the Holocene marine ­transgression brought the coastline to its present position. Marshes have spread along 18 km of the eastern coast, where the peaty floors of former ice-marginal valleys are crossed by the present coast of the Puck Bay and Puck Lagoon.

4.  Sea Level Changes and Storm Surges The most important features of sea level change on the Polish coast are as follows (> Fig. 8.8.14). The mean annual sea level rise over more than a hundred years amounts to 0.7 mm/year on the western coast (Swinoujscie), 1.1 mm/year on the middle coast (Kolobrzeg) and 1.2 mm/year in Gdansk on the eastern coast (Rotnicki and Borowka 1991).

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This tendency has accelerated markedly in recent decades. A distinct sea level rise over the years 1951–1990 has been slowest in the west (Swinoujscie 2.19 mm/year), more marked in the middle coast (Kolobrzeg 2.19 mm/year, Ustka 2.56 mm/year), and faster in the east (Gdansk 4.02 mm/year). When this is split into two 20-year intervals (1951–1970 and 1971–1990), the sea level rise is seen to have accelerated over the  last twenty years (Swinoujscie 4.38 mm/year, Ustka 4.38  mm/year and Gdansk 7.67 mm/year). A comparison of the frequency of sea level stages in the two-decade periods (1951–1970 and 1971–1990) shows a decrease in the frequency of low-water stages from 37% to 26%, no change in the frequency of average stages, and a substantial increase in the frequency of highwater stages, from 70% in Swin­oujscie to 154% in Gdansk. Among high-water stages there has been a steep increase in the frequency of storm surges which grew by 79% between the two 20-year periods. There has thus been an increase in storm surges regime in the southern Baltic (Rotnicki and Borzysz­kowska 1999), which has resulted in severe erosion on cliffed coasts and foredunes.

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⊡⊡ Fig. 8.8.13 Geomorphological sketch of the Swina Gate and Swina Spits. (Courtesy S. Musielak and K. Osadczuk.)

⊡⊡ Fig. 8.8.14 Mean annual sea level rise in the southern Baltic in the years 1951–1990 at Swinoujscie and Gdansk. (Courtesy K. Rotnicki and W. Borzyszkowska.)

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References Borowka RK (1990) The holocene development and present day morphology of the Leba Dunes, Baltic coast of Poland. In: Nordstrom KF, Psuty N, Carter B (eds) Coastal Dunes. forms and processes. Wiley, Chichester, pp 289–313 Borowka RK, Rotnicki K (1995) Shoreline changes of the Leba barrier in modern times. J Coastal Res Special Issue 22:271–274 Dobracka E, Dobracki R (1995) Geology and geodynamics of the cliff coast near between Niechorze and Trzesacz. J Coastal Res Special Issue 22:283–285 Mojski JE (ed) (1995a) Geological atlas of the Southern baltic (1:500 000). Polish State Geological Institute, Warszawa Mojski JE (1995b) Geology and evolution of the Vistula Delta and Vistula Bar. J Coastal Res Special Issue 22:141–149 Musielak S, Osadczuk K (1995) Evolution of the Swina Gate. J Coastal Res Special Issue 22:305–308

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Rotnicki K, Borowka RK(1991) Impact of a future sea-level rise in the Polish Baltic coastal zone. Rutgers University Press, Institute of Marine and Coastal Sciences, New Brunswick, New Jersey, 1–27 Rotnicki K, Borzyszkowska W (1999) Accelerated sea-level rise on the Polish coast of the Baltic. (in Polish). In: Borowka RK, Mlynarczyk Z, Wojciechowski A (eds) Ewolucja geoekosystemow nadmorskich poludniowgo baltyku. Bogucki Wydawnictwo Naukowe, Poznan, pp 141–160 Subotowicz W (1995) Lithodynamics and protection of the cliff coast at Jastrzebia Gora. J Coastal Res Special Issue 22:203–206 Tomczak A (1995) Relief, geology and evolution of the Hel Spit. J Coastal Res (Special Issue) 22:181–185 Uscinowicz S (2003) Relative sea-level changes, glacio-isostatic rebound and shoreline displacement in the Southern Baltic. Polish Geological Institute, Special Papers, 10. Warszawa, Poland, pp 1–79

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8.9 Germany

Klaus Schwarzer · Horst Sterr

1.  Introduction The coastlines of Germany have one sector facing the Baltic Sea and another facing the North Sea, separated by  the 60–80 km wide Jutland peninsula, but connected via Skagerrak and Kattegat. While the North Sea is fully marine, the Baltic Sea has a lateral gradient in salinity from Northwest (Kattegat) to East (Gulf of Finland, Bothnian Bay). Its oceanographic character is estuarine with a stratified water body, showing pronounced thermocline, halocline and pycnocline especially during the ­summer months. In the coastal areas of the North Sea and Baltic Sea wind mainly blows from SW and W, with an annual average of 35%–40%. Apart from that, winds vary around the eastern sector, i.e., from NE to SE at about 30% of the time or directly from the south at about 10%–15% of the time. The varying shape and changing orientation of the coastlines are the reason that general wind conditions are ­modified on a regional level (Lefebvre and Rosenhagen 2008). Generally speaking, the mean wind speed is lower in the Baltic Sea than in the North Sea area. Depending on the exposure of the coastline, wave action is generated mainly by winds from NW and NE–E, with patterns of longshore drift varying in relation to coastal configuration. The Southern and Southwestern parts of the Baltic Sea are mostly free of ice except during severe winter, when sea ice forms locally. Several times since 1742 the entire Baltic Sea has been covered with ice (HELCOM 2007). In the Baltic Sea the tide range is less than 0.20 m, but water level fluctuations of up to 5 m can occur during extreme storm surges, due to alternations of onshore and offshore winds, changing atmospheric pressure and seiche effects. There were major storm surges in 1864, 1867, 1872, 1904, 1913, 1949 1954 and 1995: the 1872 storm surge caused the highest water level ever measured along the  coast between Ahlbeck and Flensburg with a maximum value of 3.30 m above mean water level at LübeckTravemünde. The North Sea coast is a mesotidal to low macrotidal environment, with mean spring tidal range of 3.2 m at

Cuxhaven and 4.0 m at Wilhelmshaven. Waves are mainly created by westerly winds. The highest water level during a storm surge was measured at Cuxhaven in 1976, 5.1 m above Northern North (NN), which corresponds to MSL in Amsterdam (Jensen and Müller-Navarra 2008).

2.  The German Baltic Sea Coastline The German Baltic Sea outer coastline is about 724  km long. Small islands are connected by growing spits, initi­­ ating the development of the famous baymouth (bod­­ den)  coast  (the  term bodden is the regional name for shallow, semi-enclosed lagoons) (> Figs. 8.9.2). The length of this coastline, which is regarded as “inner coastline”, is two times longer than the outer coastline; therefore the German part of the Baltic Sea comprises about 2000 km of coastline. The Holocene history of the Baltic Sea started with the retreat of the Scandinavian ice sheet. Untill 10,300 years bp (Before Present, data are given in conventional 14-C dates) the Baltic Ice Lake developed. As result of the interaction between eustatic sea level rise and postglacial uplift of Scandinavia a marine ingression formed the Yoldia Sea, lasting for about 700 years. This was followed by another fresh water phase, the Ancylus Lake. During this time the water gradually advanced towards the present coast. Due to the global eustatic processes the Baltic Sea was again connected to the North Atlantic by way of the North Sea about 7,900 years bp (Björk 1995; Lampe 2005): the Litorina transgression. The sea rose rapidly (up to 2.5 cm/ year during the early phase of the transgression) submerging the pre-existing glacial relief of ridges and terminal moraines, undulating basal moraines and meltwater channels. About 6,000 years ago the sea level was close to its modern position. Waves and currents began to modify, initiating the development of the Bodden coast. The coastal profile by abrasion, transport and reeposition of sedi­­ ment, resulting in the development of spits, sandy hooks and beach ridges. Small islands, consisting of morainic material, were connected by growing spits, initiating the

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development of the “bodden” coast. The island of Rügen and the Fischland-Darss peninsula are examples of several small islands of Pleistocene age having been connected by beach ridges. In response to glacio-isostatic rebound, there is ongoing uplift in the northern part of the Baltic Sea, the Bothnian Bay, with rates up to 9 mm/year relative to the present sea level; in the southern part, subsidence of up to 2 mm/year occurs (Meyer and Harff 2005) intensifying erosion and coastal retreat. Along the Baltic Sea coastline cliffs alternate with ­lowlands. The cliffs are mainly cut in morainic material, and are eroding with recurrent landslides and shores strewn with boulders. However, much of the intervening coast is low-lying and beach-fringed, sometimes backed by dunes. Beach and nearshore sediment are mainly composed of sand, with some gravel and boulders, derived from cliff erosion and sea floor abrasion. Due to the decreasing distance from east to west between the terminal moraines the meltwater deposits, formed between the morainic ridges, also diminish westward. Thus sand is more abundant for beach formation in the east than in the west. The predominance of westerly winds in the South­ ern  Baltic yields a prevalence of waves from west and northwest, and consequently an eastward longshore sediment transport, along coastal sectors exposed to these directions. However, winds from the Northeast, East and

Southeast are also frequent resulting in westward sediment transport. West of Odra Lagoon the coast of Usedom Island ­consists of an alternation of cliffs, partly cut in meltwater deposits, and lowlands. Along the 40 km long outer coastline sandy beaches dominate, backed by dunes and bodden (Hoffmann et al. 2005). Cliffs and lowlands border the shores of the large Greifswalder Bodden (506 km2), SE of Rügen Island. On the NE coast of Rügen wedges of Cretaceous chalk were pushed up by glaciers (> Fig. 8.9.1), leaving some morainic material between the wedges. This cliffs are the highest along the southern Baltic coastline, attaining a maximum of 121 m above sea level at Königsstuhl, near Sassnitz. They have apron fans of chalk rubble produced by freezing and thawing in winter. These are undercut by wave action during high water levels. Recently there have been huge landslides of up to 50.000 m3. The beaches close to the chalk cliff are composed of flint shingle, with boulders resulting from glacial deposits. Rapid recession has occurred on the island of Hiddensee off the west coast of Rügen, where the Dornbusch-cliff reaches a height of 72.4 m. Spits are growing rapidly, notable the old and young Bessin in the NE of the island and the Gellen spit in the South. A large depositional feature is the Darss (> Figs. 8.9.2 and  >Fig.  8.9.3),  a major cuspate foreland built up in stages indicated by up to 105 sub-parallel beach ridges

⊡⊡ Fig. 8.9.1 Chalk cliff on the Northeast coast of Rügen. The brownish part in the cliff is boulder clay between chalk wedges. (Courtesy Geostudies.)

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formed ­during the past 3,500 years. The pattern of these ridges indicates an eastward migration of the foreland. They are truncated on the retreating western shore, while accretion continues at the Northern tip (Darsser Ort) and in Prerow Bay to the east. Bluffs of Altdarss (see > Figs. 8.9.2) to the north mark a 5,000 year old cliffed coastline, subaerially degraded behind the migrating foreland. There are several inlets, notably at RostockWarnemünde and Wismar, used as navigation channels. Erosion is extensive along this coastline, and has led to the building of wooden groins, beach nourishment and artificial dunes designed for shore protection. Along low-lying sectors, lagoons have formed where valley mouths or depressions in the glacial drift topography were submerged by the Holocene transgression and were enclosed by spits or barrier beaches. The larger lagoons are called haffs, the smaller bodden. Generally the bodden are shallow, only a few metres deep, with some narrow and deeper traversing channels. The bottom sediment are mainly silt and clay, often covered by mud with

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high organic content and peat. Due to the low salinity most of the bodden shores are fringed by reedswamp. During the past some of these low lying areas have been reclaimed for agriculture. Before the unification between Eastern and Western Germany in 1990 most of the bodden were suffering from over fertilization due to uncontrolled discharge of nutrients resulting from intensive agriculture. In the past coastal waters and shallow lagoons have been freezing in severe winters, when shore ice persisted in bays and bodden for up to 3 months. In Mecklenburg Bay and inner Lübeck Bay bodden did not develop; instead cliffed sectors alternate with lowlands fronted by nearshore sand bars (> Fig.  8.9.6). The laws of the state of Schleswig-Holstein prohibit cliff protection because these landforms are natural sediment sources supplying with sandy material. East of Lübeck the cliffs are retreating at a rate of about 0.30 m/year while west of Lübeck the average retreat is 0.24 m/year. In inner Lübeck Bay glaciers carved two deep elongated valleys. Those fjord like geomorphological landforms are given the

⊡⊡ Fig. 8.9.2 Erosion and accumulation rates of Fischland, Darss and Zingst based on the comparison of maps since 1895. Bodden have developed between the peninsula and the mainland. (Courtesy W. Jahnke.)

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⊡⊡ Fig. 8.9.3 The beach ridges of the Darss with Darsser Ort at the eastern tip. (Courtesy Walter Schumacher.)

⊡⊡ Fig. 8.9.4 Coastal evolution near Travemünde, showing the evolution of Brodten cliff. (Courtesy K. Gripp.)

local name “Förde”. Seven of these Fjord-like inlets exist between Lübeck Bay and Flensburg. They include Hemmelsdorf Fjord and Trave Fjord (> Fig.  8.9.4) separated by the up to 20 m high cliff of Brodten. This cliff has retreated about 6 km during the past 6,000 years, supplying adjacent coasts with a huge amount of sediment, blocking Hemmelsdorf Fjord completely about 1,000 years

ago. Trave Fjord is kept open artificially as the main navigation channel for the city of Lübeck-Travemünde. Slumping of soft sediment during winter and spring produces fans of shore debris that are consumed and dispersed by wave action during high water levels, distributing sand and gravel to adjacent beaches. Beach ridges indicating coastal progradation of up to 2 km are found on

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⊡⊡ Fig. 8.9.5 Niendorf harbour in the inner part of Lübeck Bay. In the background is the remnant of the former Hemmelsdorf Fjord, which is now a lake.

⊡⊡ Fig. 8.9.6 The Graswarder spit, with recurves showing stages in eastward growth. (Courtesy Office of Water Management and Rural Development, Schleswig Holstein.)

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the western and southern shores of Lubeck Bay, notably at Timmendorfer Strand (> Fig. 8.9.5) and Pelzerhaken, east of the small city of Neustadt. The west coast of Mecklenburg Bay continues to the narrow strait of Fehmarnsund. Fehmarn Island to the north has cliffs cut in glacial deposits on the east coast and beaches and spits bordering bays on the other coasts. Spits built and prolonged by longshore drifting have formed at several localities in Kiel Bay, notably Graswarder spit at Heiligenhafen, a barrier island that has grown eastward for about 3,000 years with the addition of successive recurves (> Fig. 8.9.6). In Kiel Bay are another four Fjords: Kiel Fjord, Eckernförde Fjord, Schlei and Flensburg Fjord.

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Flensburg Fjord is 20 km long and up to 25 m deep; Eckernförde Fjord and Kiel Fjord are about 15 km long and up to 20 m deep. The long, narrow Schlei inlet and Flensburg Fjord occupy valleys shaped by glacial erosion and subglacial spillways. The Fjords are bordered by cliffs and intervening lowlands. In the deeper parts mud is deposited. Erosion is predominant along the coast of Kiel Bay and low-lying sectors are protected by a variety of shore protection works. The largest is in front of Probstei at Eastern Kiel Fjord, where a dyke more than 14.3 km long and 4.5 m high is bordered by groynes with breakwaters at  their seaward ends (> Fig.  8.9.7). There is recurrent

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beach ­nourishment. Along the Baltic Sea coastline many recreational centres have been built.

3.  The German North Sea Coast The German North Sea coast includes the North and East Frisian barrier islands along the seaward side of the shallow Wadden Sea, and the funnel-shaped river-estuaries of the Eider, Elbe, Weser, Ems and Jade. The SW North Sea is shallow, the 20 m isobath being up to 50 km offshore. Tide ranges are 2–3 m in the East Frisian Islands, regions ­increasing to more than 4 m in the inner German Bight, but decreasing to 1.6 m in the N Frisian region. The semi­ diurnal tides are strong enough to generate currents that form ripples and sand waves of various dimensions (crest spacing up to 30 m) in the tidal inlets as well as in the flood and ebb deltas. At low tide extensive sand and mud­ flats  are exposed in the Wadden Sea, threaded by intri­­ cately branching channels shaped by the ebb and flow of tides. These channels vary in form, dimensions and loca­ tion as they meander and migrate laterally. The mainland coast of the Wadden Sea is protected by dykes built along the  seaward margins of marshland. The mouths of the ­estuaries are tidally scoured, with distinct ebb and flow ­channel systems. They have been modified by dredging to maintain navigation to the ports of Hamburg, Bremer­ haven, Bremen, Wilhelmshaven and Emden. Waves reaching the outer shores of the barrier islands vary in height depending on the season. The average

annual significant wave height at the North Sea Buoy II location (55° 00' N, 06° 20' E) is 1.6 m (Klein and Frohse 2008). Because of the dominance of westerly winds longshore drift is eastward along the beaches of the East Frisian Islands, while in the North Frisian Islands it is alternately northward (SW waves) and southward (NW waves). Sand washed into the tidal inlets is shaped into ebb- and flooddeltas, with shoals partly emerging at low tide. The sandy beaches are fronted by complex nearshore bar systems, shoals, swales, and channels shaped by the interaction of wave action, rip currents and tidal currents. The coastal region is subject to occasional storm surges produced by strong winds accompanying the passage of depressions across the North Sea (> Fig.  8.9.8). Storm waves during such periods accomplish substantial erosion of the sandy shores of the barrier island chain, some reshaping the intertidal morphology, and damaging the dykes that protect the hinterland. Major storm surges occurred during the years 1164, 1219, 1287, 1362, 1436, 1532, 1570, 1634, 1717, 1756, 1792, 1825, 1904, 1909, 1911, 1953, 1976, 1981, 1990 (five hurricanes at 3 days), 1999, 2000 and 2007 (Jensen and Müller-Navarra 2008). The 1362 surge inundated 1,000 km2 of land, part of which remained permanently below high tide level while the 1634 surge broke the island “Strand” into the islands Pellworm and Nordstrand, separated now by the largest tidal channel at North Frisian area, the Norderhever. In recent centuries the building, enlargement, and repair of dykes has maintained the mainland coast and much of the inner shoreline of the barrier islands, but the ⊡⊡ Fig. 8.9.7 Shore protection on the Probstei coast, outer Kiel Fjord. A dyke is combined with T-shaped groynes and breakwaters.

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⊡⊡ Fig. 8.9.8 Levels of historic storm surges recorded on a tide pole at Husum. (Courtesy Geostudies.)

outer beaches and the channels between barrier islands are subject to continuing changes, especially during stormy periods. Previously the Holocene marine transgression had in some places inundated a glacial drift topography of the Saalian ice age in the southern and Northern part of the North Sea. Sand derived from these deposits initiated the formation of barrier islands, in some cases built on or around residual foundations of glacial drift, while the finer sediment accumulated on the floor of the Wadden Sea and in the estuaries, relinquished particularly during slack water at high tides. On the seaward side of the barrier islands relatively high wave energy has built beaches of well-sorted medium to coarse sand. Although regional differences exist, partly due to the palaeo-geomorphology, the characteristic series of Holocene deposits is of fine grained sand, silt, clay and intercalated layers of peat. Within this alternating sequence, marine and brackish deposits overly­ ing peat mark phases of marine transgression. Within the Wadden Sea the mudflats possess a rich and varied benthic fauna, and the sediment are subject to

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bioturbation. Some areas are dominated by shell accu­ mulations, and shelly lag gravels occupy the floors of some of the current-scoured channels (> Fig. 8.9.9). The upper intertidal zone carries salt marsh vegetation (local name: groden), but large areas of salt marsh have been embanked and reclaimed for agriculture in recent centuries. During cold winters with NE winds ice forms on the tidal flats and occasionally on the shores of the barrier islands. South from the Danish border the islands of Sylt and Amrum consist of beaches and dunes built against a residual core of Pleistocene and Tertiary deposits. On Sylt these deposits are exposed in cliffs up to 30 m high, topped by dunes up to 20 m high (Uwe Dune, Kampen). Sylt has grown northward and southward by spit prolongation, but the west coast is straight, the glacial drift cliffs passing into cliffed dunes with blowouts (> Fig. 8.9.10), and the inner shore has marshes bordering the Wadden Sea. The entire west coast of Sylt Island is eroding (> Fig. 8.9.11). For more than a century shore protection schemes have been applied. Since 1963 mainly beach ­nourishment  has been carried out to protect the coast, partly combined with the application of geotextiles: about 37 million m3 of sand have been deposited (Hofstede 2008). At the southern end is the retreating spit, Hörnum Odde. Coastal erosion has prompted the building of jetties and tetrapod barriers, as well as artificial beach nourishment. In 1928 Sylt was connected to the mainland by the Hindenburg Dam, on either side of which there has been increased sedimentation and salt marsh formation. Amrum is similar to Sylt on a smaller scale, but in its  shelter is the rounded island of Föhr, again with a core of glacial drift bordered by salt marshes, but without major beaches and completely surrounded by man-made dykes. South of Amrum very broad intertidal sand shoals have been built up locally to about 1 m above mean high tide level by North Sea waves, and sometimes surmounted by incipient grassy dunes and beach barchans, which persist until they are overwashed and destroyed by storm surges. The broad shallow topography reduces wave action under normal conditions, and thereby protects the islands of Pellworm, and Nordstrand, each of which contains embanked marshlands. The Halligen are a group of small islets on tidal flats, totalling 23 km2 in area; they have farms sited on artificial mounds built up above mean storm surge levels. The islands have been reduced in area, particularly during storm surges, when waves generated by westerly gales cliff their shores. To the south a broad peninsula of glacial drift projects to St. Peter Ording, and

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⊡⊡ Fig. 8.9.9 Shell beds in the North Frisian Wadden Sea. (Courtesy Klaus Ricklefs.)

⊡⊡ Fig. 8.9.10 West coast of Sylt, with dunes and blowouts. (Courtesy Geostudies.)

beside it is the estuary of the Eider, restricted since 1973 by the Eider-Lock. The largest of a number of low sand shoals is the island of Trischen. West of the Elbe estuary are the low islands of Neuwerk and Scharhorn, the latter much reduced by erosion during the past century, on the broad sandy shoal that  runs out from Cuxhaven. The small island of Alte Mellum stands in a similar situation on the shoal that extends seaward between the Weser and Jade estuaries. The East Frisian Islands are a chain of elongated ­dune-capped barrier islands (Wangerooge, Spiekeroog, Langeoog, Baltrum, Nordeney, Juist, and Borkum) and three associated low sandy beach islands (Minseneroog, Memmert and Lutjehorn), extending to the Dutch border. Each has a wide sandy beach on the outer coast (>Fig. 8.9.12), backed by dunes, then a low, partly marshy inner shore. It is thought that this barrier island chain originally formed some distance to the north, and that with the secular rise in sea level in Holocene times it was driven southward by a combination of dune migration and storm surge washovers. Behind is the Wadden Sea with extensive areas of sand and mud exposed at low tide, and bordering salt marshes. The island of Helgoland, 50 km offshore, consists of red sandstone uplifted by salt dome development and an elongated sandy island, known as Düne, to the east. Waves have sculpted the margins of the sandstone into cliffs and stacks, fronted by a shore platform exposed at low tide. Much reduced in historical times by this erosion, Helgoland has also been modified by its use as a bombing range.

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⊡⊡ Fig. 8.9.11 Coastal defence at Kampen on Sylt. Protection of a short sector by artificial structures is soon outflanked by adjacent cliff recession. (Courtesy Federal Ministry for Education and Research.)

⊡⊡ Fig. 8.9.12 The North Sea beach on Norderney Island. (Courtesy Geostudies.)

References Björk S (1995) A review of the history of the Baltic Sea. Quatern Int 27:19–40 HELCOM (2007) Climate Change in the Baltic Sea Area – HELCOM Thematic Assessment 2007. Balt Sea Environ. Proc. 111:49 Hoffmann G, Lampe R, Barnasch J (2005) Postglacial evolution of coastal barriers along the West Pomeranian coast, NE Germany. Quatern Int 133–134:47–59 Hofstede J (2008) Coastal flood defence and coastal protection along the North Sea coast of Schleswig-Holstein. Die Küste 74:134–142

Jensen J, Müller-Navarra SH (2008) Storm surges on the German coast. Die Küste 74:92–124 Lampe R (2005) Lateglacial and Holocene water-level variations along the NE German Baltic Sea coast: review and new results. Quatern Int 121–136 Lefebvre C, Rosenhagen G (2008) The Climate in the North and Baltic Sea Region. Die Küste 74:45–59 Meyer M, Harff J (2005) Modelling Palaeo Coastline Changes of the Baltic Sea. J Coast Res 21(3):598–609

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8.10 Denmark

Troels Aagaard

1.  Introduction 1.1.  Geology During the Pleistocene, Denmark was repeatedly covered by glaciers. Consequently, most of the Danish coastline has been moulded in glacial or glaciofluvial deposits (>Fig.  8.10.1). Cretaceous formations are exposed in coastal cliffs in isolated patches in SE Sjælland at Stevns, on the island of Møn, and at Djursland. The coastline of the island of Bornholm is mainly formed on resistant PreCambrian and Palaeozoic rocks. The penultimate glaciation, the Saalian, covered the whole of Denmark whereas the last glaciation, the Weich­ selian, extended only to the Main Stationary Line, which runs from Bovbjerg on the North Sea coast eastward to Viborg and then in an almost straight line to the south. West and south of this line Weichselian glaciofluvial deposits are interspersed with the eroded remnants of Saalian sandy glacial and glaciofluvial deposits. As a result, the west coast of Jutland (Jylland) south of Bovbjerg is rich in sand. North and east of the Main Stationary Line, however, the coast is mainly moulded in young glacial deposits, forming sand-starved coastlines, except for areas containing raised marine deposits.

2.  Sea Level History During the Holocene deglaciation phase that started about 12,000 bp, the glaciers initially retreated from the northern parts of Denmark, and northern Jyl­land was subsequently transgressed by the Yoldia Sea. Around Frederiskhavn marine deposits from that transgression occur up to a level of about 55 m above present sea level, whereas Yoldia deposits do not appear in the southern half of the country, which was still ­ice-­covered at the time of this transgression. Isostatic rebound soon took over, and Denmark and large parts of the North Sea became subaerially exposed. A renewed transgression, the Litorina transgression, occurred around ­8,000–6,000 bp, and since then sea level has fallen in the major part of Denmark.

Because of the uneven distribution of the ice sheet thickness and forebulge collapse, the northern parts of the country exhibit the largest rebound since the Litorina transgression, up to 13 m at Frederikshavn and more at Skagen, and northern Jylland still has falling sea levels. The rebound decreases southward, and the southern parts of the country have been subjected to a continuing marine transgression. Recent sea level changes have been variable. Since 1890, the sea level at Esbjerg has been rising by some 1.2  mm/year (Nielsen and Nielsen 2002) while it is still falling by some 0.5 mm/year at Hirtshals (Christiansen et al. 1985). In combination with the differences in parent material (abundant sand in the southwest and a scarcity in the rest of the country), this gradient in relative sea level change over about the last 6,000 years has had major impacts on the development of the Danish coastline.

3.  Waves and Tides Frontal passages associated with low-pressure systems tracking along the Polar Front are frequent, especially during the autumn and winter seasons. The fronts often generate strong westerly winds and gale force winds occur for 15–20% of the year along the most exposed parts of the North Sea coast. With the dominant westerly winds and 400–600 km long fetches, the North Sea coast has a high-energy storm-dominated wave climate. The mean annual significant wave height along the central parts of this coastline is about 1.5 m. South of the Horns Rev, the mean annual significant wave height is 1.0 m. The coast between Skagen Spit and Horns Rev also has low northwesterly swell, whereas swell is practically non-existent south of Horns Rev and in the interior waters, east of Skagen Spit. The wave climate along the more sheltered coasts east and south of Skagen is also storm-dominated, but mean annual wave heights are significantly lower, on the order of 0.35 m, or less. The intricate coastline creates small and variable fetches, which result in a myriad of small-scale coastal features.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.10, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.10.1 Map showing the surface geology of Denmark. 1. Prominent coastal cliffs. 2. Glacial deposits. 3. (Glacio)fluvial deposits. 4. Marine deposits. Geographic locations named in the text are indicated. Modified from Møller (1982).

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The coasts are microtidal and consequently the coastal morphology is wave-dominated except for the area south of Horns Rev where the coasts are mixed wave-tide dominated. The tidal range along the North Sea coast decreases from about 2.0 m at the German border, to 1.6 m at Esbjerg, 0.8 m at Hvide Sande, and 0.4 m at Skagen Spit. Tides are small along the protected coasts, less than 0.3 m but significant fluctuations in water levels occur due to wind surges. Westerly winds may increase the water levels up to about 1 m above mean sea level with the reverse occurring with strong easterly winds. Surges are also important in the Wadden Sea region where water levels exceed + 2 m DNN (Danish Ordnance Datum) several times per year and + 4 m DNN about once every 10 years, on average. In the following, some typical Danish coastal stretches and sites will be described in more detail.

4. The Exposed Central West Coast of Jutland The central west coast of Jutland extends from Lodbjerg in the north, to Blåvands Huk and Horns Rev in the south. This coastline exhibits some instructive coastal changes resulting from differences in geology and process gradients in terms of wave exposure and relative sea level change. Coasts have been formed in glacial till in the north and mainly glaciofluvial deposits south of Bovbjerg; wave

⊡⊡ Fig. 8.10.2 Fallen blockhouses indicate retreat of the dune coast in Vigsø Bugt. (Courtesy Geostudies.)

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exposure decreases from north to south and the northern part has been subjected to a relative fall in sea level during the Holocene while the opposite is the case in the south. Bulbjerg is a 44 m high promontory where a steepsided Chalk ridge has been cliffed. A long gently curving sandy beach backed by eroding dunes with many tumbled blockhouses (>Fig.  8.10.2), dating from the 1940–1945 era of German occupation, runs behind Vigsø Bugt and out to Chalk bluffs towards Hantsholm. The longshore sediment transport is large. A point of divergence exists around Lodbjerg; north of Lodbjerg, the net annual drift is directed to the north while it is southward directed (except for local drift reversals at the Thyborøn barriers) and increasing to about 2.3  million m3/year at Nymindegab (KDI 2001). As a consequence of the positive drift gradient, the beaches between Lodbjerg and Nymindegab are generally eroding. The sediment is deposited between Nymindegab and Horns Rev and hence this shoreline is accreting and has prograded significantly. The coastal profiles along the west coast are very different: In the north, at Thyborøn, the shoreface profile is steep with only one longshore bar, reflecting the sand-starved status of this area. Further south, the profile slopes become gentler and the number of bars and their size increases. At Lodbjerg, the beach is backed by an about 10 m high cliff cut in Weichselian till interspersed with layers of outwash deposits and Oligocene clay floes. The shoreline is eroding at about 2 m/year and the coastal retreat and the strong winds have resulted in the breaching of cliff-top dunes

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and the formation of numerous blowouts (>Fig.  8.10.3). Behind the beach, a transgressive sand sheet has developed. The aeolian sand sheet is 5–10 km wide and reaches elevations of 32 m above sea level. The dominant dunes are parabolic. Three phases of aeolian activity have been dated by Clemmensen et al. (2001) and the most recent phase was initiated at the onset of the Little Ice Age, 1,000–1,200 ad. From historical records, it is known that major sand drift and formation of the parabolic dunes occurred 1,625–1,810 ad but the sand drift was gradually stopped by planting marram grass and establishing a forest. Indentations in this coastline have been straightened by the formation of barriers at the entrances to Limfjorden, Nissum Fjord, and Ringkøbing Fjord (fjord is a Danish term for a lagoon). The barrier beaches are wide and, as on the rest of the mainland coast south of the Main Stationary Line, they are backed by up to 25 m high foredunes. Fishery harbours are situated in the Fjords but because of the large littoral drift rates, groynes and jetties have been constructed to protect harbour entrances. The Holmsland Klit Barrier is 35 km long and 1–2 km wide. The barrier consists of several units of washover deposits (Anthony and Møller 2002). The natural inlet to the lagoon (Ringkøbing Fjord) was earlier located at Nymindegab at the southern end of the barrier. Under the  influence of the littoral drift, this inlet exhibited a

natural migration with successive cycles of downdrift ­displacement, followed by updrift breaching. This was impractical to the local fishery activities and caused problems with insufficient drainage and flooding in the hinterland. Therefore, an artificial entrance was dredged at Hvide Sande in 1910 but already in 1911 a large storm increased the width of the inlet from 26 to 230 m and the attempt to create an artificial inlet was abandoned in 1915. In 1931, a new channel was constructed and stabilised with training walls and locks and a fishery harbour was established. The channel entrance was stabilised with jetties to avoid ­silting-up of the channel, but as a consequence downdrift erosion began, amounting to about 3–4 m/year south of the inlet (>Fig. 8.10.4). Strong lee-side erosion occurs south of all harbour entrances along the central west coast. The largest erosion rates, up to 10 m/year, occur south of the large groyne field  at Bovbjerg/Fjaltring where the 40 m high coastal cliffs are protected by revetments, groynes, and breakwaters. These engineering structures, which were constructed from 1875 onwards, have succeeded in halting beach retreat at Bovbjerg, but shoreface steepening still occurs (Laustrup et al. 2000) making the coastline increasingly vulnerable. Large parts of the west coast from Thyborøn to  Hvide Sande are ­protected by engineering structures. However, realising the negative consequences of hard

⊡⊡ Fig. 8.10.3 Coastal cliff with small transgressive cliff-top dunes at Lodbjerg.

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⊡⊡ Fig. 8.10.4 Oblique air photo of the artificial entrance to Ringkøbing Fjord at Hvide Sande. The harbour jetties have interrupted the southerly directed littoral drift. Note the about 300 m downdrift offset of the shoreline. (Courtesy Niels Nielsen.)

⊡⊡ Fig. 8.10.5 The sandy beach at Vejers Strand, wide at low tide, is typical of the west Jutland coast. (Courtesy Geostudies.)

coastal ­protection, beach and shoreface nourishment was introduced in the mid-1970s and this is now the method of choice to combat coastal erosion in Denmark. Annual nourishments along the central west coast amount to about 3 million m3 per year. Erosion problems disappear south of Nymindegab where the longshore sand transport gradient is negative. Wide

beaches (>Fig.  8.10.5) and stable foredunes appear as the sand is deposited updrift of the huge natural groyne, the Horns Rev, which is probably a terminal moraine from the Saalian, extending about 40 km westward into the North Sea.  The large cuspate foreland of Blåvands Huk has developed in the lee of the shoal, and the shoreface is gently sloping with 3–4 longshore bars. Even though relative sea

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level has been and is rising in the area, this effect is overwhelmed by the sediment surplus, and a regressive barrier system has developed north of Blåvands Huk. The system started prograding south and west about 7000 bp (Clemmensen et al. 1996) and consists of sand and gravel spits that are attached to inland glacial knolls further north. The spit sequence is capped by an aeolian plain at a level of about +6 m Ordnance Datum with several foredune ridges. Further inland, parabolic dunes were active during the Little Ice Age.

⊡⊡ Fig. 8.10.6 Maps showing the successive stages in the Holocene evolution of Skagen Odde. (Courtesy Lars Henrik Nielsen, GEUS.)

5.  The Skagen Spit The Skagen Spit system forms the other end of the littoral drift along the Danish west coast. From Lodbjerg, there is a virtually uninterrupted northward longshore drift that eventually reaches Skagen where about 1.5 million m3 of sand are deposited annually (KDI 2001). The spit system has formed since the Litorina transgression. As sea level rose, it cut an up to 40 m high WNW-ESE trending steep cliff in Pleistocene glacial and late Pleistocene marine deposits at what is now the base of the spit. A spit started to grow from the western extremity of the cliff along with a succession of regressive barriers along the protected east coast of the spit (>Fig.  8.10.6). The spit system is now almost 40 km long. Because of continuing isostatic rebound, amounting to about 16 m at the root of the spit, the marine deposits reach an elevation of 14 m above sea level at the base of the spit with a northward dip of 0.67 m/km. The major part of the northerly net sand transport along the west coast is supplied by erosion of older spit deposits. This has created a steep coastal cliff with erosion rates of some 1.5 m/year. The marine deposits in the coastal cliff consist of truncated recurved spits, intercalated by peats that have developed in swales between the ridges and superimposing upper and lower shoreface sediment (>Fig. 8.10.7). Thus, the internal sedimentary structures constitute an instructive example of a forced regressive sedimentary system (Nielsen and Johnnesan 2009), and through C14-dating, the peats have allowed a reconstruction of the chronology of the spit evolution and relative sea level change (Hauerbach 1992). The end of the spit is named Grenen. Because of the shift in shoreline orientation, the longshore transport rate decreases rapidly and about 1 × 106 m3 of sand is deposited annually; the remainder of the net longshore transport is deposited on the submarine extension of the spit. Average spit progradation has been 3.8 m/year over the last 5,500

years, but large fluctuations in the progradation rate occur at decadal and centennial time scales (Hauerbach 1992). Most of the spit system is capped by aeolian sediment. The North Sea coast is and has been strongly affected by aeolian sand drift (>Fig.  8.10.8), which was particularly widespread during the Little Ice Age. The first reports of sand drift at Skagen appeared in the 1500s; roads became impassable and the church of Skagen had to be abandoned in 1795. From 1792 onwards, the sand drift was gradually halted by planting of marram grass and fir, and today stabilised parabolic dunes, some of them reaching the beaches on the eastern side of the spit, testify to the conditions that prevailed a few centuries ago. Only one transgressive dune, the Råbjerg Mile, was left in its natural state. This 800 m wide parabolic dune has advanced 3500 m inland, and its present migration rate is 12–14 m/year.

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⊡⊡ Fig. 8.10.7 The coastal cliff in raised marine deposits at Kandestederne, Skagen Odde. The cliff cuts obliquely the recurved spit ridge structures; note the peat (Martørv), which was deposited in the inter-ridge swales. (Courtesy Anders Bartholdy.)

6.  The Wadden Sea South of Blåvands Huk, the barrier chain consisting of Skallingen, Fanø, Mandø, and Rømø stretches to the German border. Contrary to the Skagen Spit, this region is, and has been, subjected to a relative sea level rise. Fur­ thermore, tide ranges are significantly larger than regions north of Blåvands Huk, making this a mixed wave-tidedominated environment. Storm surges occur frequently because of the strong westerly winds and the shallow shelf. In historic times, the most devastating surges occurred in 1362 and in 1634, when sea level allegedly reached 6.1 m above present-day mean sea level at Ribe. Since then, dykes have been erected along the entire mainland shore to protect against flooding. There is a littoral drift convergence at the barriers of Mandø and Rømø with a northerly littoral drift from the German island of Sylt, and a southerly drift from the Horns Rev. Because of the positive longshore transport gradient, the coastline of Skallingen is eroding heavily. The present-day shoreline retreat is about 3–4 m/year and the foredunes are rapidly disappearing (Aagaard et al. 1995). Skallingen probably originated as a low transgressive barrier. Early maps from 1654 and 1804 reveal a sand body almost devoid of vegetation and dunes. However, dunes began to form about 1800 with sand supplied by

­ umerous washover fans. Several dune fields accreted in n the period 1800–1900 (Aagaard et al. 2007), but sediment supplies to the barrier body were cut off in the early twentieth century, when dykes were erected by local farmers. The purpose of these early dykes was to promote establishment of marsh vegetation and improve grazing. In the 1930s, more sturdy sea dykes were built, and aeolian activity was limited to the formation of foredunes seaward of the dykes. As the shoreface retreated, accommodation space decreased. Additionally, storm surge frequency increased and the littoral drift also increased because of dredging activities in the inlet of Grådyb. Consequently, the foredunes have eroded rapidly since the 1970s and repeated surveys have shown that they are now disappearing and the dykes are under threat. Behind the barrier at Skallingen is the largest undyked salt marsh area in the Wadden Sea. This marsh started to develop in the 1930s when washover activity had stopped and deposition of fine-grained sediment started on the sandy overwash deposits. The present accretion rates are about 3 mm/year on average (Nielsen and Nielsen 2002) and it appears that marsh accretion is able to keep up with the present rate of relative sea level rise. Despite the relative sea level rise in the region, the barrier island Rømø has been accreting over a period probably spanning about 600 years due to a large ­sediment

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⊡⊡ Fig. 8.10.8 The lighthouse at Rubjerg Knude (23 m high) was built in 1900 at a distance 200 m inland from the cliff. It was abandoned in 1968 when the clifftop dune had reached a height such that the light was no longer visible from the sea. The buildings around the lighthouse have become inundated by the transgressing dune.

supply (Madsen et al. 2007). Several foredune ­systems exist, and the present progradation rate is 2–3 m/year. The beach at Rømø is extremely broad, to some extent because vehicles are allowed on the beach, which hinders the development of new foredunes. Mandø is somewhat set back from the general coastline arc of the Wadden Sea, and may be quite old, but little is known about the development of this island. It is fronted by muddy tidal flats and the sandy proto-barrier of Koresand. The development of Koresand has been traced back to the beginning of the nineteenth century (Jespersen and Rasmussen 1994), and in the context of the regional setting it is an intriguing question why dunes have never developed.

7.  The East Coast of Jutland South and east of Skagen, the coast is less exposed to strong wave action and the substrate is mainly glacial till.

The coastline is therefore intricate with numerous inlets and fjords (lagoons), and gaps between morainic hills have been filled with marine deposits. Consequently, beaches are mostly narrow and backed by till cliffs. Shore faces consist of sand-starved (till) abrasion platforms but in areas where loose sediment are abundant, a broad spectrum of various coastal landforms (recurved spit complexes, tombolos, cuspate forelands) occur because of the varying degrees and directions of wave exposure caused by short and variable fetches. Fetch length and direction are the dominant factors in shaping the coastal morphology along these coasts. The low-lying coast south of Fredrikshavn is backed by a bluff marking the Litorina Sea coastline (>Fig. 8.10.9). The eastern outlet from the Limfjørden comes to the coast east of Aalborg. At the mouth of Mariager Fjord to the south salt marshes are spreading seaward (>Fig. 8.10.10). There has been extensive diking and reclamation of tidal marshes and mudflats. The north coast of the Djursland Peninsula has slowly receding cliffs cut in Chalk with gently undulating layers of nodular flint. Associated gravel beaches contain whitecoated flint nodules and grey cobbles and pebbles derived directly from erosion of the Chalk cliffs. At Karlby Klint the cliffs rise to 11 m (>Fig. 8.10.11). At Hoed, the coast had prograded as the result of beach ridge formation beside a quarry that has dum­­ ped  large quantities on flint gravel on to the shore (>Fig. 8.10.12). There are several islands in the Kattegat, east of Jutland. They include the little island of Hjelm, which has bluffs (that were cliffs in the Litorina Sea phase) behind an emerged beach-fringed coastal lowland (>Fig. 8.10.13). The island of Lindholm (>Fig.  8.10.14) is a glacial drift hill rising to 25 m, with cliffs along its shores and spits at its extremities, one forming a linking barrier (often breached by wave scour) to the smaller eastern island of Rumpen.

8. The Coasts of Sjælland and South-East Denmark Sediment have been sourced mainly from erosion of morainic knolls. A geomorphological map of Sjællands Odde (>Fig. 8.10.15) shows the spit-like shape of the pen­ insula, owing to the connection of isolated moraine knolls by sand and gravel deposits that have been eroded from the knolls, transported alongshore and then uplifted by isostatic rebound (Schou 1949). The island of Glænø south of  Sjælland is flanked by two recurved spit complexes,

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⊡⊡ Fig. 8.10.9 The bluff marking the emerged Litorina Sea coastline is fronted by a coastal plain, the marine foreland. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.10 Salt marsh north of Mariager Fjord. (Courtesy Geostudies.)

because sediment eroded from the morainic island has been transported in both directions away from the knoll because of the (relatively) long SW and SE fetch. In some re-entrants where sediment have been more plentiful, small barriers or barrier chains have developed. Examples occur along the east coast of Jutland, in the Sejerø Bugt, Køge Bugt, at Læsø, and at Gedser. Many of the spits and barriers developed only recently, probably due to widespread dieback of Zostera vegetation in the 1930s, which laid bare expanses of sandy sediment on the

sea floor. The development of the Køge Bugt barriers has been traced in detail back to 1897 (Nielsen and Nielsen 1997) (>Fig.  8.10.16). The first thin sand bodies that started to appear in 1909 were unconnected to the mainland shore, and clearly developed from bar emergence through onshore sediment transport. The barrier system then in­­creased in size, in part through spit extension, and became attached to the mainland in 1969. There is no evidence of barrier transgression, partly because high-energy wave conditions in Køge Bugt are associated with easterly

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⊡⊡ Fig. 8.10.11 Chalk cliff at Karlby Klint. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.12 Shingle beach ridges at Hoed. (Courtesy Geostudies.)

winds that cause a lowering of Danish waters. Hence, the barriers are not overwashed when waves are large. On the north side of Fakse Bugt, the coast curves NE past Rødvig (>Fig. 8.10.17) and runs out to Stevns Klint, an east-facing chalk coast fronted by gently sloping shore platforms backed by grey shingle cobble beaches, and scattered granite boulders. There have been landslides forming aprons of talus that bear dense scrub and large blocks of chalk have fallen to the shore. The cliffs continue past a

headland at Mandehoven and diminish NW to Praesteskov, where the wooded slopes of Magleby Skov descend to the shore. At the eastern end of the island of Møn is a high ridge, consisting of Palaeocene sands and clays over Chalk. The Chalk has been displaced to the SE by ice sheets, and driven up over younger sediment. On the east coast, the cliffs (Møns Klint) attain heights of over 100 m (>Fig. 8.10.18). There have been many landslides, clay spilling down over

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⊡⊡ Fig. 8.10.13 The island of Hjelm. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.14 Cliff cut in glacial drift on the north coast of the island of Lindholm. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.15 Geomorphological map of Sjællands Odde. Brown colours indicate glacial sediment, while marine sediment infilling the gaps between morainic knolls are yellow. (From Schou 1949.)

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⊡⊡ Fig. 8.10.16 Shoreline configuration of the Køge Bugt barriers 1897–1994. Note that the barrier has accreted in place and there is no evidence of landward transgression. (Courtesy Niels Nielsen.)

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⊡⊡ Fig. 8.10.17 Chalk cliff and cobble beach at Rødvig. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.18 Møns Klint, showing apron of weathered Chalk talus. (Courtesy Geostudies.)

Chalk, often in cauldrons with a basal fan of chalky rubble. The highest part of the cliff, Store Klint, rises 128 m.

9.  Bornholm The Danish island of Bornholm stands 170 km SE of Copenhagen in the Baltic Sea south-east of southern

Sweden. It has an area of 588 sq. km and a coastline about 120 km long. The northern part of the island is dominat­ ­ed  by Pre-Cambrian granites and metamorphic rocks, ­showing a WNW-ESE trend, while Palaeozoic and some Mesozoic sandstones, shales, and limestones outcrop in the west and south. There is a capping of glacial drift, mainly boulder clay, exposed in cliffs, and some areas of sandy outwash, as at Rønne Airport.

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⊡⊡ Fig. 8.10.19 Granite gneiss forms the rocky shore at Gudhjem Harbour on the north coast of Bornholm. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.20 Gorge excavated along a diabase dyke at Helligdom­sklip­ perne, northern Bornholm. (Courtesy Geostudies.)

The north and west coasts are generally steep and cliffed, with rocky shores on granite and gneiss (>Fig.  8.10.19), and numerous slot gorges (>Fig.  8.10.20) cut out along intruded dykes of diabase, which weathers to a soft rock that can be excavated by streams. The granite and gneiss are dissected along joints in cliffs and shore outcrops. Locally, as at Helligdomsklipperne on the north coast, the granitic cliffs have been dissected into rock pinnacles. At and above high tide level, there is often a zone of black algae (>Fig. 8.10.21) and at higher levels zones of grey and orange lichen. There are segments of shore platform related to horizontal joints in the granite, some of which may have formed at a higher sea level. Much of the shore is littered with boulders, including glacial erratics. There are occasional beaches of gravel and sand, with some patches of reeds and rushes (>Fig.  8.10.22): mean salinity is about 0.8% in this part of the Baltic Sea. Some of the sandy beaches are backed by low dunes. The south and south-west coasts have sectors of cliff cut in Mesozoic rocks, notably near Arnager on the south coast, where Cretaceous chalky limestone overlies greensand in the cliffs (>Fig. 8.10.23), and south of Hasle. Shore outcrops of Cambrian sandstone are seen at Snogebaek and Nexø on the east coast. The south-east coast of Bornholm between Boderne on and Snogebaek is a large sandy lobate foreland with successively formed dune ridges bearing pine forest, and fringed by sandy beaches running out to Duedodde, the southernmost point, fronted by a broad shallow nearshore area.

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⊡⊡ Fig. 8.10.21 Slightly submerged shore platform on the NW coast of Bornholm formed by shore ice plucking. (Courtesy Geostudies.)

⊡⊡ Fig. 8.10.22 Boulders, rushes, and a gravel beach on the north coast of Bornholm. (Courtesy Geostudies.)

About 10 km north of the NE point of Bornholm are the tiny granitic Ertholmene Islands, including Christiansø; and Frederiksø; (linked by a footbridge), Graesholm, and a number of smaller skerries. The archipelago is a wildlife refuge for seabirds.

References Aagaard T, Nielsen N, Nielsen J (1995) Skallingen – origin and evolution of a barrier spit. Medd. Skallinglaboratoriet, Vol 35, Institute of Geography, University of Copenhagen, Denmark p 85

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⊡⊡ Fig. 8.10.23 Cretaceous formations exposed in the cliff west of Arnager, southern Bornholm, include a chalky Limestone over Greensand. (Courtesy Geostudies.)

Aagaard T, Orford JD, Murray A (2007) Environmental controls on coastal dune formation; Skallingen Spit, Denmark. Geomorphology 83:29–47 Anthony D, Møller I (2002) The geological architecture and development of the holmsland barrier and Ringkøbing Fjord area. Dan J Geogr 102:27–36 Christiansen C, Møller JT, Nielsen J (1985) Fluctuation in sea-level and associated coastal response: Examples from Denmark. Eiszeit Gegenw 35:89–108 Clemmensen LB, Andreassen F, Nielsen ST, Sten E (1996) The late Holocene dunefield at Vejers, Denmark: Characteristics, sand budget, and depositional dynamics. Geomorphology 17:79–98 Clemmensen LB, Pye K, Murray A, Heinemeier J (2001) Sedimentol­ogy, stratigraphy and landscape evolution of a Holocene coastal dune ­system, Lodbjerg, NW Jutland, Denmark. Sedimentology 48:3–27 Hauerbach P (1992) Skagen Odde – Skaw Spit. An area of land created between two seas. Folia Geogr Dan 20:119 Jespersen M, Rasmussen E (1994) Koresand. Die entwicklung eines ­aussensandes vor dem dänisches wattenmeer. Die Kűste 56:79–91 KDI (2001) The Danish Coastal Authority, Lemvig. Sediment budget Vestkysten. p 54

Laustrup C, Madsen HT, Knudsen SB, Sørensen P (2000) Coastal steepening in Denmark. Proceedings international coastal engineering ­conference (2000), American Society of Civil Engineers, New York, pp 2428–2438 Madsen AT, Murray AS, Andersen TJ (2007) Optical dating of dune ridges on Rømø, a barrier island in the Wadden Sea, Denmark. J Coastal Res 23:1259–1269 Møller JT (1982) Denmark. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 325–333 Nielsen LH, Johannesen PH (2009). Facies architecture and depositional processes of the Holocene-Recent accretionary forced regressive Skagen Spit system, Denmark. Sedimentology 56:935–968 Nielsen N, Nielsen J (1997) Evolution of a small, young barrier system,  Jersie Strand. Ølsebymagle Revle, Denmark. Aarhus Geosci 7:35–48 Nielsen N, Nielsen J (2002) Vertical growth of a young backbar­rier salt marsh, Skallingen, SW Denmark. J Coastal Res 18:287–299 Schou A (1949) The landscaps. Royal Danish Geographical Society (Atlas over Danmark/serie I/bd. 1/Ed.: Niels Nielsen)

8.11 The Netherlands

1.  Introduction The sandy coastline of the Netherlands borders the southern part of the North Sea between Denmark and Belgium. The total length of the Dutch coastline is somewhat more than 400 km, in which three units can be distinguished: the Rhine, Meuse and Scheldt estuary in the SW part, coastal barriers in the central part and the Wadden Islands bordering a large tidal flat area in the northern part of the country. Geologically, the Netherlands is a young country. Pleistocene surface deposits slope seaward and pass beneath Holocene deposits in the coastal area. The lowlying coast of the Netherlands, much of it below mean sea level, is protected by dunes and dykes against the sea. The present situation is the result of tectonic subsidence, a ­eustatic rise of sea level after the last glaciation, and the deposition of sand supplied by rivers or derived from the bottom of the North Sea. The tectonic subsidence is associated with the downwarping of the North Sea basin (Jelgersma 1979). During the Quaternary a large amount of sediment was deposited in the southern part of the North Sea basin. The total thickness of Quaternary sediment in the northwestern part of the Netherlands is more than 500 m, but continuing subsidence is of the order of 1.5–2.5 cm per century (Jelgersma 1961). In the early Holocene, about 10,000 years ago, the western and northern part of the Netherlands consisted of a seaward-sloping sandy Pleistocene plain. The southern North Sea was land, and Britain was still connected with the continent. A major marine transgression occurred during the Late Quaternary, and in late Boreal time the rising sea reached about 25 m below the present level as the marine transgression approached the present coastline  of the Netherlands. After that time the submerging Pleistocene land surface was gradually buried by Holocene sediment, the thickness of which varies considerably. In many places along the coast these deposits are more than 20 m thick.

The central Dutch coast shows three zones of sedimentation: in the east a zone of peat; then a middle clayey zone of tidal flats, salt marshes, and brackish lagoons with intercalated peat layers; and near the coast a littoral sandy zone of coastal barriers and dunes. The last environment is only well developed in the central part of the coastline; in the southwest and the northern part, barriers and dunes were destroyed by coastline recession after the Roman period. The presence of the clayey zone of tidal flats, salt marshes, and brackish lagoons with peat, at and below the present land surface, indicates that during the Holocene period the submerging land must have been protected from the open sea by coastal barriers or beach ridges. The present coastal area of the Netherlands is periodically subject to storms from the southwest, west, and northwest. The last direction is the most dangerous, as the configuration of the southern North Sea causes extremely high tides during NW storms. Mean spring tide range on much of the Dutch coast is 1.5–2 m (West Terschelling 1.8 m, Den Helder 1.7 m, Hook of Holland 1.7 m), but in the southwestern part it increases to 4.0 m at Westkapelle, 4.4 m at Flushing and 5.4 m at the port of Antwerp. Longshore drifting of sand from west to east is of great importance along the Wadden Islands, but less so from Den Helder southward to the Belgian border. The whole coastline can be classified as micro-to-mesotidal, depositional and wind-dominated.

2.  The Netherlands Coastline The northern coast of the Netherlands consists of a chain of sandy barrier islands (the Wadden Islands, including Schiermonnikoog, Ameland, Terschelling, Vlieland and Texel) seaward of the Waddensee, an area of intertidal sandflats and mudflats threaded by low tide channels that converge to the tidal channels between the islands. On the North Sea coast the islands have beaches backed by dunes, now eroded in many places (>Fig.  8.11.1), but

Edited version of a chapter by Saskia Jelgersma in The World’s Coastline (1985: 343–352). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.11.1 An eroding dune coast at Westerduinen on Texel Island. (Courtesy Geostudies.)

their southern parts consist of embanked and reclaimed marshland bordered by salt marshes and intertidal flats. There are indications that the coastal dunes (Younger Dunes) overlie Older Dunes and coastal barriers that formed during Subatlantic times. These are younger than the barriers in the western part of the Netherlands. The whole area of the Frisian barrier islands is highly dynamic: there is a lateral displacement of the tidal inlets between the barrier islands due to strong littoral drift. The tidal flat area of the Waddensee is rather young, having taken shape in the twelfth century when the sea invaded a lagoon and swamp area, transforming it into the present intertidal flats. South of Den Helder the coast was reclaimed from the sea after the fifteenth century. The coastline consists of dykes and young dunes, the latter partly artificial, with a core of man-made sand traps made of screens of twigs and reeds. Coastal barrier ridges are overlain by the Older Dunes, parts of which are covered by the Younger Dunes (Jelgersma et al. 1970; Van Straaten 1961, 1965; Zagwijn 1965). The coastal barriers consist of sandy ridges separated by depressions filled with peat. They formed during the Early Subboreal. The Older Dune sand accumulated on top of the coastal barriers, and the stratigraphy indicates that deposition of wind-blown sand was interrupted by periods of stability when the dunes were covered with vegetation, and phases when soil and peat beds were deposited. These phases when the formation of the Older Dunes was interrupted by vegetation colonisation have been correlated with transgression phases in the backing swamps, resulting from phases of higher precipitation rather than

⊡⊡ Fig. 8.11.2 Symbol of the Netherlands struggle against the sea. (Courtesy Geostudies.)

The Netherlands

fluctuations of sea level (Jelgersma et al. 1970). However, the alternations of peaty soils and windblown sands cannot entirely be attributed to climate and rainfall variations, as widening by accretion and narrowing by erosion of the coast caused fluctuations in the groundwater level. The formation of the Older Dunes was completed before the Roman period, when a more or less dense vegetation covered the whole dune area. The western part of the Older Dunes is covered by the Younger Dunes, deposited between the twelfth and sixteenth centuries. In contrast to the Older Dunes, the Younger Dunes are rather high, 20–50 m above sea level. Most of the Younger Dunes consist of parabolic dunes, bordered seaward by a foredune and landward by a precipitation ridge. South to Bergen am Zee the Older and Younger Dunes consist of medium-fine mainly quartzose sand, but south of Bergen the sands are much more calcareous. The mineral content is also different. Both factors cause differences in the vegetation cover, which is denser in the calcareous dunes. The coastal sands south of Bergen are largely derived from fluvial Rhine-Meuse sands deposited in the North Sea area during the Pleistocene, reworked by wave action and carried shoreward during the Late Quaternary marine transgression. During Atlantic and Early Subboreal times the topography of the submerging Pleistocene land surface included an estuary near Bergen. On the north side of this estuary, recurved spits were formed from the Pleistocene material. At the end of the Subboreal the inlet near Bergen was closed, and aeolian sands (Older Dunes and Younger Dunes) deposited. Parabolic dune forms predominate. At Ijmuiden sand accretion has widened the beach both north and south of breakwaters built to protect the  entrance to the Nordzeekanaal, which runs in to

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Amsterdam. Farther south there is evidence of another former inlet interrupting the sandy barriers, the Old Rhine inlet below Leiden. The coastline curves SW past The Hague to the Hook of Holland, which marks the beginning of the Rhine and Scheldt deltaic coast with its several estuarine outlets. For centuries this area has fought a battle against flooding by storm surges from the North Sea. Only the western side of the deltaic islands has a dune fringe; the major part of the coastline is protected by dykes against the sea. After the disastrous storm surge flooding in 1953, the so-called Delta Plan was adopted. This plan is designed to protect land and inhabitants by shortening the coastline to be defended against sea floods by closing the larger inlets and outlets with dams. After completion only the southernmost inlet giving access to the harbour of Antwerp will be left open. In 1997 a structure was completed that can be emplaced to exclude North Sea storm surges from the East Scheldt and the port of Rotterdam. On the island of Schouwen the Younger Dunes overlie  Older Dunes and coastal barriers, but in the rest of the area Younger Dunes are found on top of surface and in  the subsurface of the area suggests that coastal barriers  and dunes must have been present throughout the Holocene. After the Roman period a retreat of the coastline was accompanied by the destruction of coastal dunes and barriers. The inlets and outlets became larger, and major parts of the peaty landscape were either eroded or covered by tidal flat deposits. After the twelfth century, people began to build embankments (sea dykes) as protection against the invading sea (>Fig.  8.11.2). These have been pro­ gressively enlarged and augmented (>Figs.  8.11.3 and > 8.11.4), and form an artificial coastline (>Fig. 8.11.5).

⊡⊡ Fig. 8.11.3 Stages in the enlargement of the sea dyke at Westkapelle. (Courtesy Geostudies.)

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⊡⊡ Fig. 8.11.4 Landward slope of the sea dyke at Westkapelle. (Courtesy Geostudies.)

⊡⊡ Fig. 8.11.5 An artificial coastline: the seaward slope of the sea dyke at Westkapelle. (Courtesy Geostudies.)

References Jelgersma S (1961) Holocene sea level changes in The Netherlands. Geol Stichting Med 7:101 pp Jelgersma S (1979) Sea level changes in the North Sea basin. In: Dele E (ed) The quaternary history of the North Sea. Almquist and Wiksell, Stockholm, Sweden, pp 233–248 Jelgersma S, De Jong J, Zagwijn WH, Van Regteren Altena JF (1970) The coastal dunes of the western Netherlands; geology, vegetational history and archeology. Meded Ryks Geol Dienst 21:93–167

Van Straaten LMJU (1961) Directional effects of winds, waves and currents along the Dutch North Sea Coast. Geologie en Mijnbouw 40:333–346 and 363–391 Van Straaten LMJU (1965) Coastal barrier deposits in South- and NorthHolland, in particular in the areas around Scheveningen and Ijmuiden. Geol Stichting Med 17:41–75 Zagwijn WH (1965) Pollen-analytical correlations in the coastal-barrier deposits near The Hague (The Netherlands). Geol Stichting Med 17:83–88

8.12 Belgium

Guy De Moor · André Ozer · Irénée Heyse

1.  Introduction The coast of Belgium is about 65 km long and forms part of the sandy Southern North Sea coastline that stretches from the Schelde estuary (Netherlands) in the east to Cap Blanc Nez (France) in the west (De Moor 2006). The coastline trends about W 35° SN 55° E. Tides are semidiurnal, with a mean tidal amplitude reaching 4 m and a slight increase from west to east. In his review De Moor (2006) described the geomorphology of the Belgian coast, including the sandy

beach and coastal dunes as a natural protective belt, the Quaternary evolution of the coastal plain and recent reclamation and the nearshore system of coastal sandbanks (> Fig. 8.12.1). The anthropogenic impact of stabilizing the shore by beach nourishment was also treated. The present coast comprises three main units: a very gently sloping and fine sandy beach; a dune ridge generally less than 20 m height, with width varying from a few kilometres (at De Panne and east of Knokke) to less than a hundred metres and which is the site of numerous seaside resorts; and a backing coastal plain up to 10 km wide,

⊡⊡ Fig. 8.12.1 The Belgian Coastal Plain and the Schelde Estuary. 1 – Submerged sand banks, 2 – 1960–1980 erosive sectors, 3 – younger dunes, 4 – older dunes, 5 – outcrop of Subboreal peat, 6 – limit of the reclaimed Holocene intertidal areas, 7 – present-day intertidal flats, 8 – late Glacial aeolian sand ridge, 9 – holocene river dunes, 10 – weichselian fluvioperiglacial sands (mainly Flemish Valley infillings), 11 – uplands in Lower Pleistocene deposits, 12 – uplands in Tertiary sands and clays. BI Blankenberge; Br Bredene; H De Haan; M Middelkerke; O Oostende; P De Panne; V Veurne; KB Kwinte bank; OD Oostdyck; WB Wenduine bank; MB Middelkerke bank; OB Oostende bank; WVB Walvis bank; NB Nieuwpoort bank; SMB Smal bank. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.12, © Springer Science+Business Media B.V. 2010 (Dordrecht)

670

8.12    

Belgium

overlooked by a slightly higher land formed of Pleistocene and Tertiary deposits. This coastal plain lies at a mean level of about 2 m above low-tide, but shows a distinct microrelief of low sandy ridges (traversed by sandy creeks) and clayey depressions. It corresponds to reclaimed Holo­ cene intertidal flats, and its landward margin was limit of the Holocene marine transgression. Present-day elevation and micro-relief are mainly due to differential settling caused by drainage and reclamation. On the seaward side the foreshore declines to a broad and shallow sublittoral zone, the 5 m submarine contour running about 1,500 m in front of the low tide line. Farther offshore are the shoals and channels that form the Flemish banks. An inner row runs parallel to the coastline, an outer one lies obliquely (Codde and De Keyser 1967). The coastal plain consists of Upper Pleistocene deposits resting on Tertiary formations at a level of 25–35 m below mean sea level. The Quaternary deposits comprise an alternation of marine and continental sediment including marine Eemian formations, fluviatile and aeolian deposits dating from the last glaciation, ­mid-Holocene (Flandrian maximum) marine deposits covered by surface peat (mainly Subboreal), and the sediment of the Dunker­ quian transgressions and intermediate regressive phases, such as the older dune deposits (Paepe and Baeteman 1979). In a few places (as at Middelkerke), the shore consists of a platform cut in the Subboreal peats by wave action, proving a local subsequent retreat of at least a few kilometres. In a few other places (as at Klemskerke), the Subatlantic dune belt is backed by older dunes, indicating local progradation. The Belgian coast has been tectonically stable, at least since the last interglacial. The stratigraphic top of the Eemian intertidal flats has been found at about the same elevation as the Subboreal. A first attempt of geomorphological mapping of the microrelief of part of the Coastal Plain of the Scheldt estuary near Boekhoute was made by Heyse (1979).

2.  Coastal Evolution The evolution of the coastal plain during the Dunkerquian transgressions has been elucidated by Tavernier and his various collaborators (1954, 1970). During the Dunker­ quian II transgressions (third to eighth centuries) the whole coastal plain, outside the Schelde estuary, was inundated and became an intertidal flat situated at the back of a large offshore bar, which later developed into a barrier island complex, banked against remnants of former Holocene coastlines. This is the present-day dune belt.

Soon after the inundation, silting enabled man gradu­ ally  to reclaim the flats. Later on three main dikes were ­constructed to save the western oudland and the central oudland from inundation by the eleventh century marine transgressions. Meanwhile the rest of the coastal plain was inundated again, especially in the Yser estuary (north of Diksmuide) and the Zwin estuary (northeast of Bruges). New dikes were built to reclaim these middellands. Because of drainage and reclamation, differential settling between the low-lying sandy creeks, where the peaty subsoil had been cut, and the more elevated clay ­sedimentation flats, with peat subsoil, resulted in an inversion of the microrelief. Since the thirteenth century marine transgressions (Dunkerquian IIIB) the western Schelde estuary developed quickly, with the formation of an intertidal flat zone extending up to 20 km south of its southern bank and reaching Antwerp 65 km inland. Meanwhile, the Zwin estuary silted rapidly, cutting off the medieval harbour of Bruges from its outlet to the sea. These younger intertidal flats were gradually reclaimed behind a dense network of dikes during the following centuries. The polder deposits were studied by Ozer (1976). The Dunkerquian II and III marine transgressions had  much more influence on the coast than the older Dunkerquian I and Flandrian (which produced the Calais deposits), because the latter were much more restricted. The present extent of the Calais deposits shows that, because of the higher position of the pre-Holocene substratum, part of the coastal plain east of Ostend had only a very restricted Mid-Holocene inundation. To the west the Calais deposits are much more extensive and Holocene deposits there reach a thickness of 20 m. Nevertheless all marine Holocene sequences show a juxtaposition and succession of subtidal, intertidal, and supratidal deposits. There have also been interruptions by lagoonal deposition, especially during the formation of the surface peat. Since the eleventh century marine transgressions the Belgian coastline has mainly disturbed been only by planned inundations caused by artificial breaching of either the dune belt (Ostend, sixteenth century) or the dikes (Zwin, sixteenth century and Yser, First World War, 1914). Today large parts of the Belgian coastline are characterised by defence structures such as groins and sea walls. Artificial beach nourishment has been carried out in ­several places.

3.  The Belgian Coastline The present-day open, macrotidal, sandy beach foreshore presents a succession of runnels and ridges. It shows ­lateral

Belgium 

variations in slope, width, and sand grain size. The beach slope tends to diminish westward, while its width increases, even though the tidal amplitude is greater. In a similar way, the grain size of beach sediment decreases from the east to west. The upper foreshore has generally coarser sands than the lower foreshore, the grain size increasing westward. Near Zwin the shore sediment are less well sorted because of artificial beach nourishment (De Moor 1980). During the last few decades several sectors of the coast have been subject to more or less strong erosion resulting in a lowering of the beach and a retreat of the dune front, while adjacent sections remained nearly unchanged or even had some accretion. Such sectors extend over several kilometres and their evolution has continued over several decades. Today erosion is prominent at Knokke-Heist, Bredene-De Haan, Lombardzijde, and De Panne. These sectors are considered to be lobate forelands that formed in relation to wave and current patterns related to sea floor morphology, and eroded as a consequence of changes ­offshore. Patterns of longshore movement of coastal sands were analysed by Charlier and Auzel (1961).

   8.12

References Charlier R, Auzel M (1961) Coastal geomorphology: migration of sand on the Belgian coast. (in French). Z Geomorphol 1:5 81–184 Codde R, De Keyser L (1967) The North Sea: shore and estuary of the Escaut. (in French) Atlas de Belgique 18A 18B, Comite national de Geographie De Moor G (1980) Erosion on the Belgian coast. (in Flemish) De Aardriijkskunde, N.S. 4:279–294 De Moor G (2006) The Flemish Shore, Geomorphology and Dynamics. (in Flemish) VLIZ Oostende. 154 pp Heyse I (1979) Bijdrage tot de geomorfologische kennis van het noordoosten van Oost-Vlaanderen, Verhandeling van de Koninklijke Vlaamse Academie voor Wetenschappen, Letteren en Schone Kunsten van België, Jg. XLI, 1979, Nr. 155, 257 p., 43 fig., 6 tab., 12 diagr., 32 foto's, 22 kaarten, 1 geomorfologische kaart. (in flemish) Ozer A (1976) The morphology of polders: coastal Holocene deposits. (in French) Géomorphologie de la Belgique. Hommage au Professeur P. Macar, Liege, pp 17–27 Paepe R, Baeteman C (1979) The Belgian coastal plain during the Quaternary. Acta Univ. Uppsala., Symp. Untu Ups. Ann. Guing. Cel, 2, Uppsala, pp 143–146 Tavernier RJ, Ameryckx FS, Farasijn D (1970) Coast, dunes, polders. (in French) Atlas de Belgique, Comité national de Geographie Tavernier R, Moorman F (1954) Changes of sea level in the Flamant coastal plain during the Holocene. (in French) Geologie en Mijnbouw 16:201–206

671

8.13.1 North Coast of France

1.  Introduction The north coast of France faces the English Channel (La Manche) and has a sequence of geological outcrops that is partly similar to those of the south coast of England. The Straits of Dover originated as a valley cut through the Chalk of the North Downs and the Calais district (Artois) by the combined Rhine and Thames rivers, which flowed westward. It has been widened during phases of marine submergence by the recession of the bordering Chalk cliffs as the result of wave erosion, and perhaps during low sea level phases of the Pleistocene by periglacial action. West of Cap Blanc Nez, the Chalk gives place to underlying formations across an anticline on the coast to Boulogne and Le Treport, matching the Weald between Folkestone and Eastbourne, and to the west, between Ault and Sainte Adresse, near Le Havre, the Chalk returns to the coast forming cliffs like those of the South Downs. There is no equivalent of the Hampshire Basin syncline, however, and the Baie de La Seine is backed by Mesozoic formations equivalent to those of the Dorset coast. Palaeozoic rocks, with some intruded granites, come to the coast on the Cotentin Peninsula and in northern Brittany, matching those of Devon and Cornwall. Westerly winds and waves are dominant in the English Channel, and W to NW waves generate eastward ­longshore drifting along much of the coast, except on the western side of the Cotentin Peninsula where longshore drifting is southward (accompanied by sand blown alongshore by strong northerly winds). Mean spring tide ranges are substantial, 5.2 m at Dunkirk, 6.2 m at Calais, 8.0 m at Boulogne, 7.9 m at Etaples in the Canche estuary, 8.5 m at Dieppe, 7.7, at Fécamp, and 6.8 m at Le Havre. The unusually low tides of the Hampshire coast in England do not extend across to France, although the mean spring tide range falls to 5.4 m at Cherbourg. Tide ranges increase rapidly along the west coast of the Cotentin Peninsula to 11.5 m at Granville and 10.7 m at St Malo in the Baie of St.  Michel, and are a little smaller along the north coast of  Brittany: 9.9 m at Paimpol and 7.7 m at Roscoff. In

g­ eneral, the shore is wide at low tide, and currents are strong in estuaries, inlets, and the Bay of St. Michel.

2.  The Channel Coastline From the Belgian border westward, the coast consists of a sandy beach, wide at low tide when it is bordered by parallel sand bars and swale channels (>Fig.  8.13.1.1). The beach at Bray is backed by grassy dunes that are much eroded, and the coastline has receded, as indicated by the ruins of blockhouses that are now on the shore. At Dunkirk, a breakwater shelters the harbour beside beaches that were the site of the evacuation of Allied armies in 1940. At Calais, new dunes are developing behind a macrotidal beach (Anthony et al. 2007). To the west, a Chalk escarpment runs inland and a low cliff in grey marly Lower Chalk is fronted by bouldery rubble, and runs out to Le Petit Cap Nez, a headland of Lower Chalk. Slumping and gulleying on these retreating cliffs have been documented and illustrated (Lahouysse and Pierre 2003; Pierre 2007). This headland marks the beginning of the Opal Coast, where the sea is commonly greenish in color with suspended fine sediment. The Bay of Wissant is cut in Greensand and Weal­ den sediment, and is bordered by the dunes of Aval and Chatelet, fronting peaty marshland. There are hills of Lower Greensand and valleys in Gault before the coast rises to Cap Gris Nez, on westward-dipping Portland Lime­ ­stone, with landslides and gullies and boulder aprons on the shore. The coast turns southward, and low cliffs in Portland Limestone continue to Audresselles. There are outcrops of brown clay among the dunes and along the coast south to Wimereux, where Jurassic formations outcrop on Pointe de la Crèche. The coast then curves southwest to Boulogne, which has extensive promenades and docks. South from Boulogne are the sandy beaches of Picardie, wide at low tide, when multiple parallel sand bars are exposed, and backed by dunes that extend inland under

Edited version of a chapter by André Guilcher in The World’s Coastline (1985: 385–396). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.13.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

674

8.13.1  

North Coast of France

⊡⊡ Fig. 8.13.1.1 Dunes, beach, and bars at Bray. The arrow indicates the direction of prevailing longshore winds. (Courtesy Geostudies.)

pine forest. The sand has come mainly from the extensive Palaeocene and Eocene deposits on the sea floor to the west, with only limited contributions from the Canche, Authie, and Somme Rivers, which drain Chalk catchments. The coast is low-lying, but the Chalk rises a little way inland. The Canche estuary is funnel-shaped downstream from Etaples and its mouth is fringed by grassy dunes and a spit (Briquet 1930). The coastal plan is backed by bluffs marking the line of a Pleistocene (Monastrian) coastal cliff. It was in this region that Dubois (1924) identified the Flandrian marine transgression, which was later recognised around the world’s coasts as a major Late Pleistocene to Holocene sea level rise. South of Le Touquet and the dunes of Mayville and Berck is the Baie d’Authie, another dune-fringed estuary with wide prograding salt marshes. More dunes back the coast to the south before the wide Baie de la Somme, which has extensive sand shoals, mudflats, and salt marshes that in places have prograded rapidly in recent decades. Maximum tides may exceed 10 m here. The River Somme opens to the estuarine bay through salt marshes, and St. Valéry sur Somme stands beside a tidal channel on the southern shore. To the west is the large shingle spit, La Pointe du Hourdel, curving round into the estuary (> Fig. 8.13.1.2), with extensive sandbanks showing megaripples and scour channels. To the south, Cayeux has a long promenade behind a shingle beach and sandy shore. The Hable de l’Ault is a partly drained wetland behind a shingle ridge, bordered southward by rising slopes of Chalk, which come to the

coast at the southern end of the promenade at the little seaside resort of Ault. Here, the sandy beach gives place to a shore platform cut in Chalk backed by bluffs of frostshattered Chalk that pass laterally into vertical cliffs. The Chalk cliffs between Ault and Sainte Adresse were discussed by Prêcheur (1960). They are up to 120 m high, generally vertical, and recede intermittently by undercutting, slumping, and the ensuing removal of fallen rock debris. The cliffs are incised by occasional valleys, as at Bois de Cise, and fronted by strips of shingle beach dominated by flint nodules released from the eroding Chalk and by an intertidal shore platform. Normandy begins at the mouth of the Bresle. The Chalk cliffs and shore platforms resume to the west and continue across the mouth of the Yere valley. There are beaches of sand and gravel, particularly at valley mouths where they are derived from the Head deposits and declining cliffs in shattered Chalk. Cliff collapse at Puys has been related to seepage along impervious clay seams in the Chalk, lubricated after heavy rainfall (Dupervet et al. 2002). Dieppe stands at the mouth of the broad valley of the River Bethune, and is fronted by a wide beach of shingle and some sand accumulated updrift of a breakwater. At Le Petit Ailly, to the west, the cliffs have a capping of Tertiary sand and clay, which has been washed down the cliff face (> Fig. 8.13.1.3). The sand and clay capping contains silicified sandstone blocks known as sarsens, and this is a site where these have fallen to the shore to litter the platform. There are abrasion notches where the cliffs have been undercut by storm waves, which mobilize shingle at high

North Coast of France  

 8.13.1

⊡⊡ Fig. 8.13.1.2 La Pointe du Hourdel, a shingle spit at the mouth of the River Somme. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.1.3 Tertiary sediment washed down a Chalk cliff face at Le Petit Ailly. The boulders on the shore include sarsens (ferruginous sandstone) that have fallen from the cliff crest. (Courtesy Geostudies.)

tide, and these can lead to rock falls, with scars of freshly exposed white Chalk. In some sectors, the shingle beach is wider and higher, and acts to protect the cliff base (Costa et al. 2006). At St. Pierre en Port cliff, recession has been rapid in shattered Chalk overlain by pebble gravels and Head deposits. In this way, valley-mouth coves have been cut, as at Vaucottes, west of Fécamp (> Fig. 8.13.1.4).

Toward Etretat, the cliffs are vertical in hard silicified and almost horizontally stratified Chalk. Caves have been excavated along joints at their base (> Fig. 8.13.1.5). There are natural arches, such as La Manne Porte, and residual stacks, such as L’Aiguille (> Fig. 8.13.1.6). At Cap d’Antifer, the coast turns southwest, and cliffs capped by Tertiary clays have aprons of slumped clay along their base. The Chalk is underlain by Upper Greensand and Gault, and as

675

676

8.13.1  

North Coast of France

⊡⊡ Fig. 8.13.1.4 Valley mouth cove at Vaucottes. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.1.5 Stratified Chalk near Etretat with caves excavated along joints. (Courtesy Geostudies.)

this rises above sea level with the dip it becomes unstable, and at St.  Adresse there is a major landslide. The Lower Creta­ceous rocks outcrop along the southern side of the Seine estuary, past Le Havre. The estuary is funnel-shaped, and has a tidal bore (mascaret) that runs upriver during spring tides, but this has been much reduced by dredging and reclamation. From Honfleur, on the southern shore of the Seine estuary, the coast has low cliffs in sand and gravel, fronted by wide sandflats at low tide. At Hengueville the cliffs become sloping vegetated bluffs of clay. Sandy beaches continue past the tourist resorts of Trouville and Deauville, and at Auberville dark blocks of limestone have subsided over Oxford Clay to the shore, forming a landslide, Les Vaches Noires (> Fig. 8.13.1.7). These landslides are similar to those across the English Channel on the coasts of Dorset and the Isle of Wight, where wave energy is higher. West of the Orne estuary are cliffs and bluffs in Jurassic formations at Asnelles and landslides where limestone has collapsed over Jurassic clays that flow after saturation by heavy rain (> Fig.  8.13.1.8). The cliffs steepen in more coherent rock at Arromanches, and alternate with bluffs behind the beaches that were used for the Normandy landings in 1944. The Point du Hoc is a cliffy limestone promontory that was the site of a fierce battle. The Grand Vey is a wide inlet with broad sand and mud shoals between tidal outflow channels from the Vire and Carentan. The Lias formations are fringed along the coast by dunes and alluvium as it swings northwest to St. Vaast, then round a rocky peninsula past Harfleur, with

North Coast of France  

 8.13.1

⊡⊡ Fig. 8.13.1.6 The natural arch and sharp-pointed stack (L’Aiguille) at Etretat. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.1.7 Les Vaches Noires. (Courtesy Geostudies.)

many small headlands and inlets cut into well-jointed granite. The north coast of the Cotentin Peninsula has rocky headlands between little bays, some of which are sandy. Granite gives place westward to Devonian formations as far as Cherbourg. The port of Cherbourg backs former bays sheltered by long breakwaters, and to the west

a cliffed and rocky coast mainly in granite, continues to Cap de la Hague. There are strong tidal currents in the Raz Blanchart, the strait between Cape de la Hague and the island of Alderney. The coast turns southward along the western side of the Cotentin Peninsula and is at first cliffed and rocky,

677

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8.13.1  

North Coast of France

⊡⊡ Fig. 8.13.1.8 Landslide at Longues sur mer. (Courtesy Geostudies.)

then long gently curved beaches front the dunes of Vauville and Biville. It is cut across a series of Palaeozoic ­formations with many faults transverse to the coastline. This coast is directly exposed to westerly winds and waves (allowing for the protection of the Channel Islands), but as the tide range increases southwards wave action becomes less effective, and briefer at high tide into the Bay of St. Michel. Cap de Flamanville and Cap de Carteret are bold headlands, but sandy beaches are almost continuous and the intertidal zone widens southward. The sand has come mainly from Eocene sediment on the sea floor around the Channel Islands to the west. There has been little change on this sandy coastline during the past century. An estuarine inlet at Portbail has sandbanks with megaripples and salt marshes between tidal channels. There are similar estuarine inlets at Havre de St. Germain and near Coutances, and these and other interruptions are flanked by spits that have grown southward. South of Coutances, the hinterland is dominated by Pre-Cambrian formations. Granville is a port with a maximum tide range of about 15 m. The Bec d’Audaine shows two stages of southward growth, inner and outer, and shelters intervening salt marshes, beyond which the coast is set back with dunes behind a beach of shelly sand. To the south, this has been eroded, exposing a bench of marsh clay on either side of a tidal creek. Bluffs a short way inland are fronted by extensive saltings, then wide sandy mudflats with low tide channels that meander and migrate.

The estuaries of the See and Selune open westward into the macrotidal Bay of St. Michel, with Mont St. Michel a  prominent high hill rising from the coastal plain (> Fig. 8.13.1.9). The salt marshes on the southern shores of the Bay of St.  Michel have shown phases of severe ­erosion, with marginal cliffing, dissection, and the widening of tidal channels (> Fig. 8.13.1.10). Storm surges have emplaced cheniers of shelly sand on the marshes. The Pre-Cambrian rocks and associated granites of western Normandy extend into the Brittany peninsula, where they are interspersed with various Palaeozoic formations, including volcanic intrusions. At Cancale, the coast rises to the cliffs of Pointe du Grouin, the beginning of the steep coast of northern Brittany. Pre-Cambrian ­formations outcrop along the south coast of the Gulf of St. Malo, and to the west is a sequence of rock formations similar to those of western Cotentin. Slope-over-wall coast profiles are extensive along this north coast, the slope being a mantle of periglacial Head while the wall is a cliff of exposed hard rock. There has often been little erosion on these cliffs in Holocene times, waves having done little more than sweep away the lower part of the Head mantle to exhume cliffs that were shaped much earlier, in Pleistocene phases of relatively high sea level. The rocky shores are irregular, locally with segments of abrasion shore platform. There are numerous valley-mouth rias, but the wider bays generally have almost flat sandy beaches, some backed by dunes that have often been damaged by vehicle traffic and camping.

North Coast of France  

 8.13.1

⊡⊡ Fig. 8.13.1.9 Mont St Michel, with salt marsh. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.1.10 Dissected Chapel Marsh, Bay of St Michel. (Courtesy Geostudies.)

St. Malo stands close to the estuary of the Rance River, which has been dammed to generate tidal hydroelectricity. The coast then runs out to Cap Frehel a promontory with a cliffed headland of gently dipping red Permian sandstones. Cap d’Erquy has been quarried, and to the west schists and gneisses dominate the cliffs and bluffs that run behind the Bay of St. Brieuc. The Pre-Cambrian formations here include some basic volcanic rocks. The wide

intertidal flats have splays of tangue, a coherent calcareous silt formed of broken shells of local origin. The coast then trends northwest to Paimpol, which has a dissected plateau hinterland, and out to granitic Pointe de l’Arcouest. Further north, the large sand and shingle spit, Sillon de Talbert, diverges from the coastline as it swings southwest. It has been supplied with sand and shingle by eastward longshore drifting, and piled up by westerly storm waves.

679

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8.13.1  

North Coast of France

Its seaward flank has large rounded cobbles and boulders, while the inner shore has sand and pebbles, descending to a narrow tidal bay with patches of salt marsh. Roscoff stands on a granite peninsula, fronted by an up to 2 km wide irregular rocky and gravelly shore exposed at low tide. These wide rocky shores are the outcome of several phases of marine erosion during Pleistocene marine transgressions, alternating with episodes of periglaciation, which produced Head deposits that have

been partly removed during and since the Flandrian marine transgression in Holocene times. Some of the rocky stacks have trails of sand and gravel derived from Head deposits and emplaced by converging waves in their lee, as well as by rising tides (Guilcher et al. 1959). Beyond Plouescat, the dunes of Keremma back a bay, which merges into the Grève de Goulven to the west. Low cliffs of Head back bays at Plougerneau, and a wide rocky intertidal shore extends out to the lighthouse at Lilia

⊡⊡ Fig. 8.13.1.11 Wide rocky intertidal area at Lilia, looking out to La Vierge lighthouse. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.1.12 The ria at Aber Wrac’h. (Courtesy Geostudies.)

North Coast of France  

(> Fig. 8.13.1.11). Aber Wrac’h (> Fig. 8.13.1.12) and Aber Benoit are major rias opening northwest, and the steep rocky coast curves southward to Pointe de Corsen. A chain of small rocky islands runs out to the Isle of Ushant, girt by an irregular coast of hard rocky cliffs, mainly granites. There are erratic rocks on Ushant and the Molène Archipelago. Some are of distant derivation, and may have been delivered on sea ice or in icebergs. These include young basalts (dated 2.2–4.4 my) that came from Iceland or the mid-Atlantic ridge. Ushant is here taken as marking the beginning of the west coast of France.

References Anthony EJ, Van Lee S, Ruz ME (2007) Embryo dune development on a large, actively accreting macrotidal beach, Calais. Earth Surf Process Landf 32:631–636

 8.13.1

Briquet A (1930) Le littoral du Nord de la France et son évolution morphologique. Libraire A. Colin, Paris Costa S, Hénaff A, Lageat Y (2006) The gravel beaches of north-west France and their contribution to the dynamic of the coastal cliff-shore platform system. Z Geomorphol, Supplementband 144:199–214 Dubois G (1924) Recherches sur les terrains quaternaires du Nord de la France. Memoirs du Sociètè Géologique du Nord 8:1–356 Duperret A, Genter A, Mortimore RN, Delacourtt B, De Pomerai MR (2002) Coastal rock cliff erosion by collapse at Puys, France. J Coastal Res 18:52–61 Guilcher A, Adrian B, Blanquart A (1959) Les queues de comète de galets  et de blocs sur les cotes nord-ouest de la Bretagne. Norois 6:125–145 Lahousse P, Pierre G (2003) The retreat of Chalk cliffs at Cap Blanc-Nez (France). J Coastal Res 19:431–440 Pierre G (2007) Duration of the marine evolution and Holocene retreat of a cliffed coast: a case study in northern Boulonnais. Quaternaire 18:219–231 Prêcheur C (1960) Le littoral de la Manche de Sainte-Adresse à Ault, Etude morphologique. SFIL Poitiers

681

8.13.2 West Coast of France

1.  Introduction From the Island of Ushant, the Brittany coast runs south to the large Rade de Brest ria and the wider Bay of Douarnenez. The Armorican massif has folding and faulting along mainly west to east lines, and these structures influence the outlines of peninsulas and bays along the southern coast of  Brittany (Guilcher 1948). The Bay of Audierne faces SW  into Atlantic Ocean swell, and the southern coast of Brittany is embayed with rocky headlands such as Quiberon, outlying island such as Ile de Groix and Belle Ile, and inlets of which the Gulf of Morbihan is the largest and most complex. At St Nazaire is the estuary of the Loire, and the coast swings southward along the Bay of Biscay with large bays and promontories, mainly of Mesozoic rock, and several islands: Noirmoutier, Yeu, Ré; and Oléron. The wide Gironde estuary is followed by a long, almost straight sandy coastline, backed by dunes and interrupted by the Bassin d’Arcachon. In the far south, this curves out to Biarritz and the cliffs cut across the strongly folded Mesozoic rocks of the Pyrenean foothills extend to the Spanish border. Westerly winds and waves are dominant in the Bay of Biscay, with Atlantic Ocean swell arriving from the southwest and northwest. These waves generate some eastward longshore drifting on the south coast of Brittany, and southward longshore drifting on the east coast of the Bay of Biscay. Mean spring tide ranges are large in western Brittany, 5.7 m at Le Conquet, and 6.1 m at Brest, but diminish along the southern coast to 4.4 m at Concarneau and 4.7 m at Le Croisic. Much of the east coast of the Bay of Biscay has mesotidal ranges: 4.7 m at Les Sables d’Olonne and 4.5 m at Pointe de Grave on the southern side of the Gironde estuary. These tides are sufficient to generate strong currents in estuaries and straits, and to expose substantial intertidal areas of sand, mud, and salt marsh as the tide ebbs.

2.  The Atlantic Coastline South of Pointe de Corsen, the coast is steep and rocky, with slopes mantled with periglacial Head deposits and

cliffs in Palaeozoic rock. Beaches of calcareous sand occupy coves. The coast curves westward, and the wide Plage des Blancs Sablons, backed by grassy dunes, faces northwest on the northern side of the Kermorvan Peninsula. Cliffs and coves have been cut into slaty rocks (>Fig. 8.13.2.1), and on the southern side is the long estuary of Le Conquet, which has eroding salt marshes at its head (>Fig. 8.13.2.2). To the south is a steep coast to Pointe de St Mathieu, which marks the entrance to the Goulet de Brest. The steep coast, interrupted by sandy coves, becomes more sheltered as the strait narrows through the Goulet de Brest. This opens into the wide Rade de Brest, a major ria with the city and port on Brest on its north coast. The long, straight Elorn estuary comes in from Landerneau, and several inlets converge in the southern arm. The bordering coasts are hilly, with low promontories and coves with gravelly beaches and several spits, notably the shingle Sillon des Anglais (>Fig. 8.13.2.3) (Guilcher et al. 1957). To the south is the Crozon Peninsula, with hard Ordovician sandstones outcropping on an irregular stormy western coast around Camaret and rocky cliffs south to Cap de la Chèvre. The large Baie de Douarnenez is a little more sheltered, but dominated by cliffy coasts and small sandy coves, which continue round the Pointe du Raz. The wide Baie d’Audierne is cut in weathered mica schist between headlands of harder crystalline rock. It is backed by a long gently curving shingle barrier beach, shaped by refracted Atlantic Ocean swell and retreating landward in successive storms, which produce overwash fans. The bordering headlands have been little modified since the Holocene marine transgression, because they retain segments of Pleistocene emerged beaches. From Pointe de Penmarch, the south-facing cliffed coast runs eastward to successive bays with estuaries at Pont l’Abbé and the mouth of the Odel at Bénodet. There is a curved shingle barrier to the east, fronting a lagoon that has an indented inner shore. The beach curves out to a cuspate point at Mousterlin, and continues to Beg Meil, where the broad Baie de la Foret is inset. The coast remains steep and rocky east to Lorient, where there is a barrier spit beside the mouth of a ria. Offshore is the Ile de Groix,

Edited version of a chapter by André Guilcher in The World’s Coastline (1985: 385–396). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.13.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8.13.2

West Coast of France

⊡⊡ Fig. 8.13.2.1 Cove at Pointe de Kermorvan. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.2.2 Salt marsh creek at Le Conquet. (Courtesy Geostudies.)

mainly of Pre-Cambrian schist and gneiss, which has broad terraces of marine planation (>Fig. 8.13.2.4), slope-­ over-wall coasts with a mantle of periglacial Head partly removed by wave action to expose rocky cliffs, shore platforms cut across local structures, sandy coves, some with veneers of reddish mineral sand, clefts cut along joints, and a lobate sandy foreland at the eastern end. There is a similar foreland at the eastern end of Belle Ile, both the outcome of longshore drifting along the north and south coasts and accretion to the lee.

The coast curves southeast to the rocky Presqu’ile de Quiberon, which is attached to the mainland by the dunecapped tombolo of Penthièvre. The western coastline has receded as the result of erosion. To the east, a series of narrow headlands and inset bays extends past Carnac to the Pointe de Kerpenhir at the narrow entrance to the large Golfe du Morbihan. This is a marine embayment of intricate configuration, with many bays, headlands, and islands, and is thought to be partly the result of tectonic subsidence extending the area of submergence of a number of valleys:

West Coast of France

8.13.2

⊡⊡ Fig. 8.13.2.3 The Sillon des Anglais, ashingle spit bordering the Rade de Brest. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.2.4 Shore platform cut across Pre-Cambrian schist, Ile de Groix. (Courtesy Geostudies.)

Auray, Marle, and Noyalo. There is a partly ­submerged megalithic monument on the island at Er Lannic. The bordering shores have gravelly beaches and salt marshes, and wide areas of gravel, sand, mud, and seaweed are exposed at low tide. Tidal currents are strong at the entrance, but tide range within the Gulf of Morbihan is smaller than on the outer coast. East from Port Navalo at the mouth of the Gulf of Morbihan, the coast has slope-over-wall topography (>Fig. 8.13.2.5), cliffed and rocky with sectors of sandy

beach. It runs in beside the funnel-shaped bay at the mouth of the Vilaine River, where there are extensive intertidal flats, and to the south the cliffs culminate in the granite headland of Le Croisic. On the northern side, the spit at Pen Bron shelters the lagoon of Guéande, with extensive salt marshes, and with another barrier to the east forms a double tombolo linking the rocky Batz peninsula. The south-facing coast at La Baule has a series of beaches that have received sand from the Loire estuary. The Loire transports shoals of sand downstream and into

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8.13.2

West Coast of France

⊡⊡ Fig. 8.13.2.5 Slope-over-wall coast south of the Golfe du Morbihan. (Courtesy Geostudies.)

shoals off its mouth. It has a sediment load of about 0.5 million tonnes/year, delivered mainly during floods, but the downstream movement of sand and silt has been modified by dredging. Sandy sediment is carried to beaches, while plumes of mud are delivered to salt marshes, notably to the south. The Vendée coast to the south has hilly promontories and the broad shallow Baie de la Bourgneuf, sheltered by the elongated Ile de Noirmoutier. Land reclamation has been extensive here, and parts of the coast are walled. The west coast of Ile de Noirmoutier is more exposed, and there are low cliffs and ledges of Cretaceous sandstone and limestone on the coast to the south, where a broad platform of weed-strewn planed-off strata extends in front of the Maison de Clemenceau. Les Sables de l’Olonne stands at the southern end of a lagoon separated from the sea by a forested dune barrier partly overlying rocky terrain. A wide curving sandy beach, shaped by refracted southwest swell, fronts the resort promenade, and to the south of Le Bac are low rocky cliffs trenched by gravelly gullies cut out along joints, Le Puits d’Enfer. The embayed cliffy coast to the SE becomes more sheltered as it passes into the Pertuis Breton in the lee of the Ile de Ré. It culminates in a spit at the Pointe d’Arcay, which has multiple dune ridges and ends in eastward recurves. It is followed by a second spit, the Pointe de l’Aiguillon, which shelters the Anse de l’Aiguillon, a bay with a broad intertidal mudflats on to which salt marsh in spreading. The Ile de Ré is low-lying with sandy beaches, particularly on the exposed SW shore, some dunes up to 20 m high, salt marshes in the northern bays, and low rocky

headlands. It lies off the port of La Rochelle, with the bay, Pertuis d’Antioch, to the south. There is a broad sandy beach at Chatelaillon and the hilly Fouras promontory, which runs out to the narrow Pointe de la Fumée and outlying Ile d’Aix. The coast is sheltered by the Ile d’Oléron, and has intertidal mudflats off the mouth of the Charente River. Changes have been rapid around the Charente ­estuary, with growth of spits and rapid expansion of marshes (Regrain 1980). The Ile d’Oléron has low cliffs, beaches of sand, and shingle, and some sectors stabilised by sea walls. There is evidence that sea level stood higher here within the Holocene (Allard et al. 2008). To the south, a sandy beach is interrupted by the narrow Pertuis de Maumusson, then resumes to end in the Pointe de la Coubre (Castaing and Jouanneau 1976; Facon 1965). This is a sandy foreland with dune ridges indicating southward growth and phases of westward progradation, the sand having drifted southward along the shore. Pointe de la Coubre marks the northern side of the Gironde estuary. The Dronne, Lalinde, and Garonne Rivers converge near Bordeaux, and bring down a large sediment load, about 2.5 million tonnes/year, which circulates in the estuary with tidal ebb and flow and the effects of the Coriolis force (Allen 1973) before discharging into the Bay of Biscay. Point de Grave, south of the Gironde mouth, marks the beginning of a long sandy coastline, which runs southwest, then almost due south. The sand has come largely from the Gironde, and drifted southward, but there is also calcareous sand, sometimes shelly, swept in from the floor

West Coast of France

of the Bay of Biscay. There is a nearshore fringe of crecentic bars (Castelle et al. 2007). The beach is backed by afforested dunes on a wide barrier that encloses lakes, the Etang d’Hourtin/Etang de Carcans, and the Etang de Lacanau,in a north–south corridor of marshland. The lakes have patchy fringes of reedswamp. The sandy beach ends southward in the recurved spit at Cap Ferret, where historical maps show intermittent growth since 1708, punctuated by a phase of truncation in the mid-nineteenth century. The spit stands on the seaward side of the Bassin d’Arcachon, a triangular marine inlet fed by several small rivers, with a mean spring tide range of about 4 m. Growth of the spit has caused southward migration of the entrance, which is encumbered by sand shoals, and the advance of Cap Ferret has been matched by the scouring back of the sandy shore at Pyla to the southeast. Peat is exposed at the base of the sandy cliff (>Fig. 8.13.2.6), and the recession of the shore has maintained instability in the backing dunes, which form the Pyla Dune, an escarpment about 100 m high of wind-blown sand spilling inland over pine forests (>Fig. 8.13.2.7 ). South of Arcachon, the Landes of Gascony coast is backed by a fringe of formerly transgressive dune ridges 60–70 m high and up to 6 km wide. The ridges have been stabilised by pine plantations (introduced from 1801 onwards by Nicolas Brémontier) except for the still-active Dune of Pilat (Pyla), a ridge of bare sand 118 m high, with an escarpment spilling on to forested dunes. The coastal

⊡⊡ Fig. 8.13.2.6 Peat exposed at base of sand cliff, Pyla. (Courtesy Geostudies.)

8.13.2

dune fringe is retained by a sparse grassy and shrubby cover, but to landward the dunes are under pine forest. They backed by a hinterland sandy plain that rises gradually inland to more than 40 m above sea level, and bears heath and pine plantations. Pinchemel (1980) commented that “the vast sand covering of the Landes is believed to be an aeolian deposit laid down during a Quaternary [Pleistocene] cold spell. The material probably came from the continental shelf, where Pliocene sands are extensive on the sea floor, were uncovered by the sea during a phase of marine regression”. At Pointe d’Arcachon, the sandy coast resumes its southward trend, and the backing dunes become lower and more stable, though still with many blowouts. Minor outflow of streams is maintained at Mimizan and Contis, and the Courant d’Ouché is deflected 4 km southward along Moliets beach behind a longshore spit. The coastal dunes here are retained by a sparse grassy and shrubby cover, but those to landward are under pine forest. There are further outflows from dune lakes at Vieux-Boucau and Hossegor and the mouth of the Adour River at Boucau. Breakwaters built to protect such outflows interrupt the southward longshore drifting, and cause accretion to the north and erosion to the south. Along much of the 230 km coast between Pointe de Grave and the mouth of the Adour River, there has been coastline recession, in places attaining 10–20 m/year. Sand continues to arrive on the shore, but is lost landward to dunes, and it is possible

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8.13.2

West Coast of France

⊡⊡ Fig. 8.13.2.7 The crest of the Pyla dune, spilling landward. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.2.8 Cliffs in strongly folded Cretaceous rocks, Basque coast south of St Jean de Luz. (Courtesy Geostudies.)

that the transverse profile of the sea floor has been migrating landward. At Biarritz, the dunes come to an end as bluffs of shale and occasional cliffs of sandstone rise behind the beach. To the south, the cliffs become higher and bolder, cut in stratified Cretaceous limestone and sandstone. St Jean de Luz backs a bay cut in shale between sandstone headlands

at the mouth of the Nivelle valley, with breakwaters built to shelter the harbour. To the west, the Côte Basque has cliffs and coves cut in the strongly folded and faulted Cretaceous limestones of the Pyrenean fringe (>Fig. 8.13.2.8). These come to an end in the beach bordering a spit that has grown partly across the Hendaye bay, where the Bidassoa River reaches the sea. To the west is > Spain.

West Coast of France

References Allard J, Chaumillon E, Poirier C (2008) Evidence of former Holocene sea level in the Marennes-Oléron Bay. Comptes Rendus Geoscience 340:306–314 Castaing P, Jouanneau JM (1976) Les mécanismes de formation de la flèche de la Coubre. Bulletin Institut Geologique, Bassin d’Aquitaine 19:197–208 Caastelle B, Bonneton P, Depuis H (2007) Double bar beach dynamics on the high energy French Aquitaine coast: a review. Marine Geol 245:141–159

8.13.2

Facon R (1965) La pointe de la Coubre, étude morphologique. Norois 12:165–180 Guilcher A (1948) Le relief de la Bretagne meridionale. H. Potier, La Roche sur Yon Guilcher A, Moign A (1977) Coastal conservation and coastal studies in France. Geograph J 143:378–392 Guilcher A, Vallantin P, Angrand JP, Galloy P (1957) Les cordons littoraux de la rade de Brest. Bulletin Comité Océanographique et Etude des Côtes 9(1):21–54 Pinchemel P (1980) France. Translated by D. and T. Elkins Regrain R (1980) Géographie physique et télédétection des marais charentais. University of Amiens

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8.13.3 Mediterranean France

1.  Introduction The French coast on the Mediterranean seaboard is divided  into three sectors: a short, mountainous, jagged coast bordering the Pyrenees; a longer coastal plain in Roussillon and Languedoc, ending in the Rhône delta; and another mountainous sector in Provence and Nicois. It is a coast with a very small tide range, and wave action generated mainly by southwesterly and southerly winds, which produces longshore drifting from west to east on beaches. The climate is Mediterranean, Marseilles having mean monthly temperatures ranging from 17°C in January to 29°C in July and an annual rainfall of 640 mm, with a winter maximum.

2.  The Mediterranean Coastline From the Spanish border cliffs cut in the Cambrian and Ordovician formations of the eastern Pyrenees

(>Fig. 8.13.3.1) decline towards the Quaternary coastal plain around Perpignan. This coastal plain of Roussillon and Languedoc has a barrier coastline with low dunes enclosing lagoons that were former bays closed in the late Holocene (Martin 1978). They include the Etang de Canet, the Etang de Leucate, which has inner and outer barriers built successively, the Etang de Lapalme and the Etang de Gruissan. The long gently-curving coastline is unusual in that it has not been shaped by ocean swell. It is receding, possibly because of diminished sand supply from the Rhône River. There are promontories caused by three isolated volcanic or calcareous hills from which the barriers have grown. The Hérault River comes to the coast over a deltaic plain south of Agde, and to the east is the promontory of Cap d’Agde. The Etang de Thau, located in a syncline, is the only lagoon still having an outlet to the sea at the port of Sète. Long-term sedimentary processes have been documented by Schmidt et al. (2007). More barriers with low dunes fronting lagoons fringe the coast south of Montpellier on the western side of the Rhône delta.

⊡⊡ Fig. 8.13.3.1 Cliffs near Cape Cerbère, Spanish border. (Courtesy Geostudies.)

Edited version of a chapter by André Guilcher in The World’s Coastline (1985: 385–396). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.13.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8.13.3

Mediterranean France

⊡⊡ Fig. 8.13.3.2 Calanque, Port Miou. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.3.3 Promenade des Anglais, Nice, showing shingle beach. (Courtesy Geostudies.)

The Rhône delta is one of the most interesting in the world in spite of its moderate size, because it clearly shows the results of the shifting course of the distributaries. Two forces are in action: the mistral, a powerful northwest wind with a very short fetch, and the marine a more moderate onshore southeast wind with a much longer fetch. The strongest wave action thus comes from the southeast (Guilcher 1958). In 1710 the main distributary of the Rhône was the Bras de Fer, which was then abandoned for

the Grand Rhône at the same time as the Petit Rhône discharge was decreasing. As a result two former deltaic lobes related to these distributaries were actively eroded and their sediment transferred westward to Beauduc and Espiguette points by longshore drifting caused by the southeasterly swell, while a new spit fed by the Grand Rhône sediment (84% of the solid discharge) grew across the Gulf of Fos. More recently, however, river outflow from the Rhône catchment has been harnessed by many large

Mediterranean France

dams, so that the sedimentary discharge to the sea decreased from 40 million tonnes/year at the end of the nineteenth century to 12 million tonnes in 1956, and 4 or 5 million in 1970. Erosion largely prevails now except at Gracieuse Point where, as at Beauduc and Espiguette, growth continues locally and moderately with a reworking of the previous sediment supply. Between the Rhône delta and the Italian border, rocky coasts prevail, except at such places as the double tombolo at Glens near Toulon. Cretaceous limestones form cliffs in the Marseille district and east towards Cassis. The famous calanques east of Marseilles (>Fig. 8.13.3.2) are drowned valleys cut into Cretaceous limestone. The bordering cliffs show basal solution notches. The many interglacial beaches and associated calcareous aeolianites show that recent erosion is generally insignificant, but there has been erosion of beaches in coves in Provence. Near Bandolo the Cretaceous formations recede behind Jurassic outcrops along the coast, and in the Toulon region the cliffs are cut into Permo-Triassic formations. Relative sea level changes have been documented at Frejus (Devillers et al. 2007). East of the coastal lowland at Hyères the irregular coastline is cut across Lower Palaeozoic and PreCambrian formations that strike generally north-south.

8.13.3

Beyond Cannes these give place to Triassic, Jurassic and Cretaceous formations of the French Alps. The coastal scenery is impressive, but the Côte d’Azur from Cannes to Menton is covered by residential houses and thus completely artificial. At Nice the promenade is fronted by a shingle beach derived from gravel brought down by the River Var (>Fig. 8.13.3.3). East of Menton the Cretaceous rocks pass beneath Eocene formations that outcrop along the steep coast across the Italian border.

References Devillers B, Excoffon P, Morhange C (2007) Relative sea level changes and coastal evolution at Forum Julii, Frejus. C R Geosci 339:329–336 Guilcher A (1958) Travaux récents sur le delta du Rhône. Annales de Géographie 67:156–159 Guilcher A, Le Demezel M (1988) Man’s response to coastal changes in France. In: Ruddle K, Morgan WB, Pfafflin JR (eds) The coastal zone: Man’s response to change. pp 91–126 Martin R (1978) Evolution de deux lagunes du Roussillon depuis le maximum marin holocene. Bulletin d’Association Francaise Etude Quaternaire 14:108–111 Schmidt J, Jouanneau JM, Weber O (2007) Sedimentary processes in the Thau lagoon from seasonal to century time scales. Estuar Coastal Shelf Sci 72:534–542

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8.13.4 Corsica

1.  Introduction

2.  The Corsican Coastline

Corsica is a mountainous island with numerous peaks over 2,000 m, rising to 2,710 m on Monte Cinto. The geological map shows the western part of the island dominated by a dissected crystalline massif of various granites, with an area of acid volcanic formations in the northwest. A corridor of Miocene limestone runs NE from Bonifacio in the far south, and spurs of the same limestone project across the Plain of Aleria in the east. North of Bastia the coast steepens alongside a peninsula of  Jurassic and Cretaceous rocks with some Lower Palaeozoic volcanic formations that extends northward to Cap Corse. The west coast, made of Variscan granites, has a topography produced by marine submergence in which cliff retreat has been very slow. The major ridges trend NE–SW, and the irregular coastline includes rocky promontories and many inlets and coves aligned in this direction. Pleistocene (Tyrrhenian) beaches and dune calcarenites, capped by angular periglacial debris formed during the Last Glacial phase of the Pleistocene period, are found in a very large number of coves (Conchon 1980; Ottmann 1958). The east coast is quite different, lower and much flatter, consisting ­generally of large amounts of Pliocene and Pleistocene marine and continental deposits resting on Alpine metamorphic schists. Lagoons and barriers occur, and there has been neotectonic activity. As usual around the Mediterranean, the coastal Quaternary deposits are well preserved, largely because of hardening by calcium carbonate precipitation. Tide ranges are small (c. 0.6 m), and wave action is generated mainly by westerly and northeasterly winds across the bordering seas. Wave energy is generally stronger on the west coast than from the Tyrrhenian Sea. The climate is of Mediterranean type, with a dry summer and a rainy (often stormy) winter, and forest and scrub (maquis) vegetation is extensive.

Ajaccio, the capital, stands on the southern side of a high peninsula on the mid-west coast, looking across the gulf at  the mouth of the Gravona valley. At the head of the Gulf of Valinco is Propriano (> Fig. 8.13.4.1), where the Rizzanèse River valley opens to a marshy lagoon behind a steep beach of coarse granitic sand. Because of the small tide range the intertidal zone is narrow, and dunes are poorly developed. At Roccopina Late Quaternary marine submergence of steep-sided valleys has produced long promontories and inlets, bordered by boulder-strewn granite slopes. Rivers descend to small salt marshes and minor beaches. To the south the Montagne de Cagna has splintered outcrops and exfoliating domes of Oberonic granite. To the southeast the granite gives place to limestone, which forms the plateau of Bonifacio, terminating seaward in high cliffs (> Fig. 8.13.4.2) trenched by steep-sided inlets (calanques). Irregular cliff recession has isolated a number of stacks and islands off Capo Bertusato, which faces Sardinia across the narrow (11 km) Strait of Bonifacio. On the southeast coast a valley descends to Porto Vecchio and the Stabiacco estuary opens to a gulf of irregular outline, with bay-head salt marshes, and shoals, beaches and spits shaped by local wave action. To the north the mountains descend steeply to the coast at Solenzara, where there is a cobble beach derived from fluvially-supplied gravels, but there are also sectors of sandy beach backed by low dunes and some cliffy sectors (> Fig.  8.13.4.3). The mountain front is then fronted by the scrubby coastal Plain of Aleria, which consists of several sandy and gravelly terraces incised by valleys and truncated by interfluvial cliffs. Several lagoons occur at valley mouths between Solenzara and Bastia the Etang d’Urbino and the Etang de Diana on either side of the deltaic lobe of the Tavignano River. The lack of marsh encroachment in these lagoons may be an indication of continuing subsidence (Tricart 1954). The Etang de Diana

Edited version of a chapter by André Guilcher in The World’s Coastline (1985: 385–396). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.13.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8.13. 4

Corsica

⊡⊡ Fig. 8.13.4.1 Sandy beach at Propriano. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.4.2 Limestone cliffs and stacks at Bonifacio. (Courtesy Geostudies.)

is bordered by bluffs. As the Plain of Aleria ­narrows northward, a beach fringes low cliffs and some dunes at Prunete and there are dune segments, as at Moriani Plage. The Etang de Biguglia has a broad outer sand barrier backed by a spit that may be the remains of an earlier, older barrier. From Bastia a steep coast bordering a high ridge runs north to Cape Corse, and on the western side is the Gulf of Saint Florent. The modern beach is backed by an emerged

beach, overlain by dune calcarenite and red gravelly loams of periglacial origin (Ottman 1958). Small bays and headlands line the coast of Balagne, with a larger promontory at  Calvi. Interfluvial ridges from the rugged Tartagine Mountains descend to form headlands between valleymouth gulfs. Sandy beaches back the bays, as at Monetta, and locally sand has been swept inland by strong winds. South of Calvi the coast is generally steep, cut into ­outcrops

Corsica

8.13. 4

⊡⊡ Fig. 8.13.4.3 Cliffs north of Solenzara. (Courtesy Geostudies.)

⊡⊡ Fig. 8.13.4.4 Steep coast on crystalline rocks near Porto. (Courtesy Geostudies.)

of rugged bare granulite, as at Porto (> Fig. 8.13.4.4). At Olmo the gravelly Marsalino River has a dry watercourse with boulders and gravel in summer, but runoff after heavy winter rains has supplied grey pebbles to a stormpiled shingle ridge at its mouth. Sandy beaches are found

in more ­sheltered coves at the mouths of smaller streams, as along the hilly coast south of Cargèse, where the headlands have rocky shores. There are sandy beaches in bays at Cap  de Feno, the sand extending offshore on the sea floor (> Fig. 8.13.4.5).

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Corsica

⊡⊡ Fig. 8.13.4.5 Bay beaches on Cap de Feno, NW of Ajaccio. (Courtesy Geostudies.)

References Conchon O (1980) Les niveaux quaternaires marins et la tectonique en Corse. Niveaux marins et tectonique quaternaires dans faire mediterranéenne. University of Paris I, pp 271–282

Ottmann F (1958) Les formations pliocenes et quaternaires sur le littoral corse. Memoir, Societé; Géologique de la France Tricart J (1954) Ecological evidence of subsidence in Corsican coastal Lagoons. Revue de Géomorphologie Dynamique 5:15–19

8.14 Spain

Eric Bird

1.  Introduction

2.  The North Coast of Spain

Spain covers most (85%) of the Iberian Peninsula and has  a coastline 4,750 long. Much of the country is high plateau, the average altitude being about 700 m, and the coast is generally steep and rocky, with cliffs and bluffs, and very few low-lying areas (Marques and Julia 1985). There is a contrast between the Atlantic (Bay of Biscay and southwest Spain) coast and that bordering the Med­ iterranean, related to prevailing wave energy regimes and tide ranges. The Bay of Biscay and southwest coasts have strong wave energy, including ocean swell of distant origin, frequent storms, substantial mean spring tide ranges (San Sebastian 3.8, Santander 3.7, La Coruna 3.3, and Cadiz 2.7 m) generating strong tidal currents, and rivers with a large and constant discharge. The Mediterranean is characterised by low to moderate wave energy, small tide ranges (diminishing east from Gibraltar 0.9 m to typically less than 0.2 m), weak tidal currents, and rivers with low, intermittent discharge, with the exception of the River Ebro, the largest Spanish river flowing into the Med­ iterranean (drainage basin 85,700 sq km). Thus while steep, cliffed and rocky erosional predominate on the Atlantic coast, with large estuaries and rias, the Med­ iterranean coast is made up of large bay head beaches between rocky headlands and small deltas (the largest is the Ebro delta, 388 sq km), and narrow coastal plains fronting mountain ranges. The climate is of Mediterranean type, with higher rainfall on the northwest coast and the influence of African aridity in the south. La Coruna in the northwest coast has mean monthly temperatures ranging from 10.5°C in January to 19°C in July and an average annual rainfall of 800 mm, while Barcelona on the Mediterranean coast has 8°C in January and 23.5°C in July and an average annual rainfall of 525 mm. Apart from ocean swell from the Atlantic, wave regimes are determined by the winds from the NW to SW prevail along the northern and southwestern coasts, while easterly winds are common along the Mediterranean seaboard.

The north-facing coast of Spain, bordering the Bay of Biscay (Cantabric Sea) runs close to the northern limit of the Iberian tectonic plate, which is located at the edge of the continental shelf a short distance offshore. Much of the coastline is rocky and cliffed, with bay head beaches, small re-entrants at the mouths of valleys drained by short rivers and larger gulfs, known as rias, formed by marine submergence of deeply incised wide valleys. The coastal slopes of dipping flysch-type deposits have been undercut by marine erosion to form steep cliffs that are sometimes vertical. These are backed by terraces at various levels, locally called rasas (Guilcher 1974), cut across geological structures by marine and subaerial erosion and influenced by tectonic movements. On the border between France and Spain the estuary of the Bidasoa River opens into a large bay, with steep bluffs extending past Hondarrubia to the harbour at the river mouth. The cliffed coast begins here, with a lighthouse on the headland at Cabo Higuer. To the west is the steep coast of the Jaizkibel Range, where the coastal slopes generally follow seaward dipping yellow-brown sandstones and shales in a structure that trends east-west, parallel to and south of the mountain chain of the Pyr­enees. There are flatiron structures on the seaward-dipping rocks, and valleys cut into sharp synclines, and the coast is trenched by many small valley-mouth inlets between sharp-edged rock ridges and precipitous dip-slope cliffs. Jaizkibel rises 455 m above sea level, and the high ridge runs parallel to the coast until it declines to the Oiartzun valley. This river opens to a wide estuary below Lazo, with the old town of Pasaia on the northern side of a bay that narrows to a long, deep outlet. The steep coast continues westward to San Sebastian (Donostia), which stands behind two sandy beaches in curved bays separated by the rocky promontory of Monte Urgull. The sharply curved western beach, La Concha, is also protected by mid-bay Santa Clara Island. To the west the cliffs rise steeply to Monte Igeldo at the start of another coast-parallel high ridge. The coast is again steep, with

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.14, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.14.1 The renourished beach at Deba. (Courtesy Geostudies.)

yellow sandstone and clay formations dipping directly or obliquely seaward. The narrow road had been frequently repaired after cracks parallel to the coastal slope, which was unstable, with many slump scars where the sandstones are underlain by grey clay. The ridge ends at the mouth of the Rio de Orio valley, which runs from south to north in a transverse valley through the coast ranges. At the river mouth there are breakwaters and a wide sandy beach backed by an esplanade. The cliff on the eastern side is a dip-slope cliff, while the western side has folded yellow and grey strata culminating in a bold escarpment cliff. At the base of the cliffs the sandstones had disintegrated into an apron of boulders with no shore platform. The steep coast resumes, cut across a hilly ridge of hard red sandstone along to another valley, where Zarautz, at the mouth of a river has a seafront sandy beach. Beyond this the highway N634 becomes a corniche road, with concrete walls down to a bouldery shore, with rocky ribs and gravel but no sand. Fences have been built to intercept rocks falling from the undulating strata in the roadside cliff. A sandy beach then curves out into the tombolo, which attaches the high island of Getaria to the mainland. On the sides of the island rock strata dip steeply, and become almost vertical along the coast. To the west the corniche road continues, with boulders and occasional ramps of seaward-dipping sandstone at the base of the sea wall, many rock falls and some slumping, towards Zumala. This town stands beside a bay, the Rio Ibara flowing out between large breakwaters, and on the eastern side there is a wide sandy beach.

The river mouth is sandy but there is a salt marsh upstream, where sloping banks of mud are exposed at low tide. Westward the steep dip-slope coast descends to segments of sawn-off seaward dipping strata on the shore. Debeka Hondartza is a bold promontory east of Deba, which stands in a bay at the mouth of the Ego River valley. The sandy beach in the bay at Deba was renourished in 2001, when 215,000 m3 of sand were pumped ashore and shaped to a beach profile by bulldozers (> Fig. 8.14.1). The steep coast is interrupted by another bay at Mutriku, where there is a small sandy beach beside a breakwater. Rainfall increases westward, and the coastal slopes are well wooded. The bay at the mouth of the Artibay River at Ondarroa is largely enclosed by massive breakwaters forming outer and inner harbours. Beyond Atxazpi the coast road runs along a bridge fronting a steep cliff with horizontal strata, and there are shore segments of shore platform on gently dipping sandstone south-east of Leketio. San Nicolas is a large limestone island with a notch along the base of the cliff. The steep coast, interrupted by valley-mouth inlets at Ea and Elantxobe, has segments of limestone cliff with basal notches and stacks as at Ogono Makurra and Ogono Lurmuturra (> Fig. 8.14.2). The wide Ibaia estuary has a threshold of inwashed sand, forming broad sand spits with no dunes, and washover lobes emplaced by Bay of Biscay storm surges. This has been a UNESCO reserve since 1984. Upstream the estuary becomes muddy, with patchy salt marshes, as the river narrows towards Gernika. West of the estuary is Bermeo, then the long peninsula descending to Matxitxako.

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⊡⊡ Fig. 8.14.2 Limestone cliff on the Cantabrian coast at Ogono Lurmuturra. (Courtesy Geostudies.)

The steep coast, with many small bays, then runs on to Villano Lurmuturra and turns southwest, past the Gorlizko Hondartza estuary to the headland at Galea and the wide gulf of Abra where the Nerbioi estuary widens downstream from Bilbao. To the west the steep, cliffy and embayed coast extends past Castro Urdiales to Laredo, where a long curving sandy beach borders the triangular lowland east of the Ria de Treto estuary. Santona stands on its northern shore, and a cliffy coast curves round to Punta del Aguila. The northfacing coast is then fairly straight, curving out to points such as Cabo de Ajo, and broken by numerous inlets, of which the Bahia de Santander is the largest, fed by the Miera and Pisueha Rivers. Rio de Mogro, Rio de San Martin de la Arena, Rio de San Vicente, and Rio de Tina­ mayor are valley-mouth inlets, and to the west is the Costa Verde with small headlands and bays, extending past Llanes. The hinterland is dominated by a sequence of terraces (termed rasas) standing about 50, 120–145, 180–220 and 240–260 m above sea level in front of the escarpment of the Cantabrian Mountains (Guilcher 1974). The Asturias coast is generally steep and partly cliffed, with some segments parallel to the strike of formations in coastal ranges and others cutting obliquely across them. There are many valley-mouth inlets, notably the Ria de Villaviciosa. Gijon is a port in a bay, and an irregular cliffed coast runs out on a triangular foreland to Cabo de Peñas. To the west the Ria de Aviles and Ria de Pravia are elongated estuaries, and the coast truncates successive south to north mountain ranges, one of which culminates in Cabo

Vidio. Beyond Cabo Busto the coastline is set back southward, and continues with minor irregularities past Ria de Navia and Ria del Eo. The coast then swings northwest to Cabo Burela, and  becomes more indented, with long promontories between deep valley-mouth rias: the Ria de Vivero, Rio de Sor, and Rio de San Marta de Ortigueira. The longest of the promontories runs out to Punta de la Estaca de Bares. Vixia de Herberia is the highest cliff in Europe (613 m) stands at the end of a Sierra range. Cliffs descend to shore platforms, the present one forming as the result of the dissection and lowering of a similar Pleistocene shore platform that existed a metre or two above the present level (Trenhaile et al. 1999). At Cabo Ortegal the coast turns southwest and becomes relatively straight, apart from the small ria at Cedeira. Between Cabo Prior and Cabo Priorino the coast is embayed, and then it is broken by the three large rias, Ria de Ferrol entering from the east, Ria de Betanzos from the southeast and Ria de la Coruna from the south. Although essentially similar to the rias along the north coast, these are the first of the large Galician rias. Although essentially drowned valleys, they are considered disproportionally large for the small rivers that drain into them, and are thought to have been influenced by tectonic subsidence of continental shelf sectors, causing them to be enlarged by valley deepening (Cotton 1956). La Coruña, a long high breakwater, protects the harbour, which stands beside a broad ria, with many promontories and inlets. To the west a succession of bays and

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⊡⊡ Fig. 8.14.3 The granite promontory at Punta Roncudo. (Courtesy Geostudies.)

headlands extends to Punta de las Olas, beyond which is  the fishing village of Caión, on a low rocky promontory. To the west the beach becomes a sandy barrier, the Playa de Baldaio, a coarse quartzose beach surmounted by low dunes and enclosing a shallow lagoon with shelly deposits and salt marshes. At its mouth a rocky promontory borders an outlet through a derelict sluice. The lagoon is divided by embankments, some incorporating segments of earlier dune barriers, into large ponds used for shellfish production. Behind, rural slopes rise to forested ridges. Porrilado has a long beach with dunes and cliffs in bedded yellow and red ferruginous Tertiary sandstone over gnarled grey quartz-veined granitic rock. This is one of several small Tertiary basins along the north coast of Spain. At Razo there is a wide sandy beach exposed at low tide between cliffs of periglacial Head over rocky metamorphosed sandstone. Malpica de Bergantiños has a large harbour in a deep cove, with inner and outer breakwaters beside a high northern sea wall and rampart on which waves break heavily. Former cliffs beside and behind the cove have been stabilised by stone walls, but on the western side there has been a large landslide, and some buildings have been lost. A steep slope-over-wall coast runs westward to Cabo de San Adrián and the outlying high rocky Islas Sisargas. At Cabazo a sandy cove backed by sparse grassy dunes is bordered by a rocky shore where at low tide jointguided shore platforms are exposed. A terrace about 5 m

above high tide level is seen on headlands and a nearby flat-topped island. A terraced valley descends to a small sandy beach at Praia Barizo. To the southwest a wild storm-lashed coast extends to Punta Roncudo, a rocky promontory in coarse porphyritic granite with large micas (> Fig. 8.14.3). A steep slope rises to tors, and on the southern side a more sheltered coast has low benches and shore platforms guided by gently inclined joint planes. This marks the northern coast of the Ria de Corme y Laxe. The village of Corme, in a bay to the east, stands behind a small sandy cove and a wide rocky shore exposed at low tide (> Fig. 8.14.4), fringed by brown seaweed and kelp. There are cliffs cut in Head, and the sandy beach is backed by a grassy dune in front of an old cliff. Atlantic swell moves into the Ria de Corme y Laxe, fitting curving beaches of inwashed sand. Towards the head of the ria is a sandy cove, Praia di Balarés, then a wide sandy barrier spit, with dunes blown up the adjacent hillside. This barrier spit is the outcome of a contest between Atlantic waves moving into the ria, the ebb and flow of tides, and outflow from the river. It stands in front of an estuarine lagoon at the mouth of the Rio Anllons below Pontecesco, where there are extensive salt marshes backed by reedswamp, and including some areas that have been reclaimed, then abandoned. These incorporate remnants of an earlier inner barrier. On the southern coast of the Ria de Corme y Laxe a valley opens northward to Laxe, with a bay in the shelter of another cape, Punta Insúa. A curving beach of pale

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⊡⊡ Fig. 8.14.4 The rocky shore on jointed granite near Corme. (Courtesy Geostudies.)

⊡⊡ Fig. 8.14.5 Cabo Vilan. (Courtesy Geostudies.)

grey quartz sand with mica has a wide berm backed by uncliffed hummocks of grassy dunes with scattered marram clumps and an open Agropyron sward. This stands beside an esplanade and a small harbour behind a stone breakwater. Southwest from Punta Insúa the steep coast is broken by a wide bay containing the Playa de Traba. Atlantic swell has shaped a gently curving sandy beach back by grassy dunes and the Lagoa de Traba, with a gently rising

hinterland. On its southern side cliffs and steep slopes in granite run along to Camelle, a valley-mouth harbour, and Arou, a rocky cove with a boulder beach above firm sand exposed at low tide. Cabo Veo and Cabo Tosto have shore platforms related to almost horizontal joint planes in the granite. Vegetated slopes on periglacial rubble (Head) descend to shores of blocky granite, and are interrupted by deep joint-guided clefts. The coast runs out to the granite promontory of Cabo Vilan (> Fig. 8.14.5).

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To the south is the broad Ria de Camariñas, bordered by bays and headlands, and narrowing to the mouth of the Rio Grande. The town of Camariñas is in a sheltered bay on the northern side. Upstream, of extensive mudflats are exposed in front of rushy salt marshes, up to Ponte do Porto. On the southern side is Muxia, a small harbour in the lee of a narrow granite peninsula that runs out to Punta de la Barca. There is a curving sandy beach, Piagge de Lourido, from which dunes have moved up the hillside. A ridge runs out to Cabo de la Boutra, and to the south the coast curves round to Cabo Touriñan, where a hilly landscape declines to a windswept granite promontory. This is the westernmost point in Galicia (9° 18’ W). To the south the steep rocky coast is interrupted by curving sandy beaches, the Playa de Frixe and Playa de Rostro, shaped by refracted Atlantic ocean swell. Beyond these, the bold Costa da Mortes and the high Cabo de la Nave dominate the coast towards the south-projecting Finisterre Peninsula. Cabo Finisterre (Fisterra) has clearly-defined slopes of bouldery Head descending to low rocky cliffs and shore outcrops, an example of the Atlantic coast slope-over-wall association that reaches its southern limit near Lisbon in Portugal. The narrow peninsula has a sandy isthmus behind Finisterre, an old town behind a harbour to the northeast, and shelters a wide bay with headlands between curving sandy beaches backed by dunes, as at Sardiñeiro de Abaixo. Redonda is a peninsula on the northern shore, with a steep grassy and bracken-covered coastal slope on ⊡⊡ Fig. 8.14.6 Cabo de Corrubedo. (Courtesy Geostudies.)

Head deposits and a small basal cliff of Head over metamorphic rock. On its eastern side is a ria with Corcubion, a small town and harbour, then Brens (Cée) with a carbide factory beside an inlet with a beach of inwashed white sand and low grassy dunes at the ria head. To the south-east the steep-sided Xallas valley descends to Ezaro, where water pipes descend a steep slope to a hydroelectric station, behind a dune and beach. There are bare rocky islands out in the bay. South of Pindo the coast passes beyond the shelter of Cabo Finisterre, and wave energy increases. There are storm-eroded rocky shores and beaches of shelly sand. The Playa de Carnota occupies a large west-facing bay, and is a sandy barrier beach capped by sparsely grassed dunes and backed by a rushy lagoon. To the south the rocky coast runs out to the Punta de los Remedios, and there are coastal slopes mantled by periglacial Head deposits. The coast swings southeast past the Playa de Lariño and Playa de Louro beaches to Punta Carreiro, a rocky headland which marks the beginning of a wide gulf, the Ria de Muros y Noia, narrowing to the mouth of the Ria Tambre. On the north coast of the Ria de Muros y Noia the hillfoot town of Muros stands beside a bay. The ria coast has sandy, gravelly and rocky embayed shores, meagre dunes and muddy sediment and marshes only at the heads of inlets. Atlantic waves diminish into the ria, but storm surges occur, and there was flooding at Abelleira in 1994. Rio Tambre enters the ria head below Ponte Nafonso through an area of tussocky grass marshes and reedswamp. Noia stands beside an inlet to the south, also with

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grassy and reedy marshes. The southern coast of the Ria de Muros y Noia has a succession of small sandy and shelly beaches separated by steep rocky sectors. Wave energy increases southwest of Porto de Son, where spurs of granite protrude at Queiruga and Caamaño, and long sandy surf beaches such as Playa Rio Sieira are backed by sparsely grassed dunes and small lagoons. To the southwest is the granite promontory of Cabo de Corrubedo (> Fig. 8.14.6). The embayed granite shores of Cabo Corrubedo have large rounded boulders, spurs and knobs of granite, and patches of earthy gravel Head. There is a broad zone above high tide swept clear of soil and vegetation by frequent Atlantic storm waves. On the eastern side of the promontory is the town of Corrubedo, behind a small bay, and then the long sandy Playa de Ladeira, a southwest facing barrier beach backed by uncut grassy dunes and a reedy lagoon. Behind the grassy dunes is a large bare elongated dune, with lobes of sand spilling landward. The Sierra de Barbanza range runs southward to a promontory, Punta de Couzo, on the western side of the large Ria de Arousa. Aguina and Castiñeiras are small harbours with boulder breakwaters on the embayed Ria de Arousa coast, where rocky peninsulas occur between coves of shelly sand and gravel, as at Palmeira, where granite ribs cross the shore. In contrast with the rias to the north, there are many rocky islands, and the gulf narrows to the mouth of the Rio Ulla. The large irregular beach-fringed Illa de Arousa is linked to the mainland near Vilanova by a causeway, and the coast to the south is intricate, with an isthmus linking Illa da Toxa having a long surf beach, Playa de Lanzada, on its western shore. Mountainous peninsulas run southwest on either side of the Ria de Pontevedra, and marine submergence gas left a chain of islands offshore, continuing the alignment of the coast range to the south. Beyond the high Peninsula de Morrazo is the deep and wide Ria de Vigo, with the large port of Vigo on its southern side. An irregular coastline extends down to the high Monte Ferro peninsula and the bay of Baiona, and on out to Cape Silleiro. Here the coastline straightens southward along the western flanks of a high coast range to Santa Tecla and the estuary of the Rio Minho, much narrower than the Galician rias. The Portuguese boundary runs down the Rio Minho to the coast (> Portugal).

3.  The South-West Coast of Spain From the Portuguese border at the mouth of the Rio Guadiana the coast of the Gulf of Cadiz is dominated by a

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long gently curving sandy beach extending round to the mouth of the Rio Guadalquivir. Salt marshes and tidal mudflats with meandering creeks are extensive behind sand spits in inlets southeast of Ayamonte, but otherwise sandy beaches are almost continuous (Playa de Castilla), interrupted by the El Rompido outflow, deflected eastward by the Punta della Gato spit, formed by longshore drifting: it grew 4 km eastward in a century (1873–1974). Another interruption is the Odiel-Tinto estuary below Huelva, de­­ flected eastward by longshore drifting to form the spit at Punta Umbria. Terraces end in gullied sandy cliffs behind a 200 m wide prograded coastal fringe at Playa de Mazagon, the inner half vegetated and the outer half a sandy beach. The coastal lowlands correspond with the Neogenic depression of the Rio Guadalquivir, which is situated between Sierra Morena and the Betic Cordillera. Behind the coastal sand barrier and dune fringe there are marshes (Las Marismas) which extend up to 50 km inland, drained by channels that flow to the Guadiana, Tinto-Odiel, and Guadalquivir rivers. Beach ridges are numerous, indicating stages in lagoon, marsh, and dune evolution, the dunes becoming larger to the southeast, where they are drifting in over vegetated swales. Part of the wetland lies within the Doñana National Park. South of the mouth of the Rio Guadalquivir a hilly coast extends from Punta del Camaron to Punta Candor, at the entrance to the Bay of Cadiz. A sandy barrier spit links an island on which the port city of Cadiz stands, and within the sheltered by the Rio Guadalete has built a large delta. There are extensive salt marshes and some lagoons (salinas) used as saltworks. To the south the hilly coast continues to Cabo Trafalgar and Tarifa Point, with rocky shores alternating with sandy beaches. The coast then swings ENE, cutting across the Sierra de Ojen range on the northern shore of the Strait of Gibraltar, to the broad Bahia de Algeciras and the high limestone promontory of > Gibraltar.

4.  The Mediterranean Coast of Spain Much of the Mediterranean coastline follows the strike of  geological formations in the several sierras. Bays cor­ respond with Neogenic tectonic depressions and steep and rocky sectors with the mountain chains. The coastal fringe is tectonically active, especially in the southern part. Erosion is prevalent on many of the beaches, and structures including groynes and nearshore breakwaters have been built to conserve a beach at many seaside resorts. East from Gibraltar the Costa del Sol is backed by mountain ranges, drained by numerous sub-parallel valleys

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with streams that have supplied grey sand to beaches. There are seaside resorts at Estepona and Marbella, and the blunt delta of the Rio Guadalhorce. At Malaga the coastline turns eastward and is generally steep and mountainous, broken by valley mouths with small deltas, such as that of the Rio Adro. East of Adro there is a coastal lowland, Llanos del Almeria, then the Gulf of Almeria, into which the Rio Andorax has built a cuspate delta. The coast then runs out to Cabo de Gata. The eastern coast of Almeria trends NE and is steep alongside the Sierra del Cabo de Gata and Sierra Cabrera. The Rio Almazoro drains to a deltaic lowland, but beyond Aguilas the coast truncates the foothills of the Sierra del  Almenara. A series of headlands (Cabo Cope, Cabo Tiñoso) and bays then extends past Cartagena and out to Cabo de Palos. Small areas of Quaternary deposits spread out in front of the Sierras and produce beaches that vary in width according to the structure of the coast and the intermittent streams that deliver sand and gravel to the shore. To the north is the Costa Blanca. Mar Menor is a large coastal lagoon separated from the sea by a narrow sandy barrier with an outlet near the northern end. A generally low-lying coast continues north to Torrevieja, which has two brackish lagoons, Laguna Salada de Torrevieja and Laguna Salada de la Mata, separated by a low ridge north of the town and linked by canals to the sea. A lobate lowland runs out to Cabo de Santa Pola, with Isla Plana a Nueva Tabarca outlying. The Bay of Alicante is backed by a coastal lowland, and beyond the port of Alicante the land runs out to Cabo de las Huertas. The coastal plain then narrows in front of mountain ranges to Benidorm, where the Peñas de Arati forms a coastal ridge along a promontory. An embayed coast with cliffed promontories then extends northeast to Cabo de la Nao, where the coast swings NW. Intricate bays separate Cabo Negro, Cabo de San Martin and Cabo de San Antonio at the end of the Sierra de Montego ridges. Beyond Denia the coastal plain widens past Oliva along the Gulf of Valencia, and becomes swampy. Cap Cullera marks the beginning of a sandy barrier coast with dunes fronting swamps and the large Albufera lagoon. The Turia River delta borders Valencia and to the north to coast trends NNE with many sandy beaches known as platjas. The Rio Palancia and Rio Mijares have built small deltas on a narrow coastal plain backed by mountain fronts, the Sierra d’Oropesa and Sierra d’Irto forming steep and cliffy sectors up to Peñiscola. Here a limestone island has been attached to the mainland by a sandy isthmus, forming a tombolo. The coastal plain varies in width, widening at the mouths of valleys such as that of the Rio Cerbol at Vinaroz.

The coastline then passes behind the large southern spit of the Ebro delta. At San Carlos de la Rapita the Ebro (Ebre) delta begins. Its inner boundary is marked by segments of low limestone scarp with coves and caves, the coastline that predated the growth of the delta. Of cuspate form, the Ebro delta has large trailing spits on its southern and northern flanks. The southern spit consists of a narrow isthmus, El Trabucador, where the sandy beach has a sparse shell cover, and is backed by low dunes and an irregular lagoon shore, where some low sectors have been overwashed. This curves round to the west, past the lighthouse on Punta de la Baña, and a widening salt marsh on its northern side includes shallow lagoons and extensive salinas with heaps of salt, out to the end of the spit, Punta de las Alfaques. The southeast shore of the delta is beach-fringed, cut  across former distributary channels and intervening lagoons and marshes. There is a wide beach of firm light brown sand at Platya dels Eucaliptus, where a former distributary outlet closed in 1749. The present mouth of the Ebro River is bordered by extensive Phragmites reedswamp and salt marsh (with Salicornia, Limonium, Salsola and Juncus), with areas of dried and cracked mud deposited by river floods. A dune ridge runs NW from the river bank, behind the reedfringed Bassa del Gorxal lagoon, which is a faunal reserve with many seabirds including pink flamingoes. An outer barrier with low dunes carrying tamarisk, eryngium and bare sand patches encloses a lagoon on Cap de Tortosa, and curves round to the seaside town of Riumar. Here the coast has prograded, and a wooden walkway leads out from the esplanade across a sandy swale (the end of the lagoon zone) to outer dunes, where wind-drifted sand is being intercepted by fences to form dunes stabilised with marram grass and shrubs (> Fig. 8.14.7). The annual sand yield from the Ebro, formerly about 3 million tonnes, has dim­inished because of dam construction upstream, but the beach at Riumar continues to prograde (like Malindi in Kenya and Seaside in Oregon on this is a seaside resorts disadvantaged by continuing progradation). Longshore drifting to the northwest has supplied sand to the northern (Fangar) spit, but sectors of this sandy coast have been retreating, as at Platja de la Marquesa, where an embankment has been built to protect a restaurant that now protrudes from the receding beach. The Fangar spit continues to grow, sheltering aquaculture in the Port del Fangar bay. The delta plain is crossed by levees and reedy areas, which indicate former distributary channels, and is a landscape of ricefields drains and irrigation channels, with some reed-fringed delta lakes, such as Bassa de l’Estella. To the north wave energy increases in the Golfo de

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⊡⊡ Fig. 8.14.7 Prograded beach at Riumar. (Courtesy Geostudies.)

⊡⊡ Fig. 8.14.8 Nearshore breakwaters have shaped the beach at Sitges. (Courtesy Geostudies.)

San Jorge, beyond the shelter of the Ebro delta. Between La Ampolla and Salou a narrow coastal plain is backed by a gently-rising coastal pediment crossed by dry river channels descending from limestone mountains. Beaches show longshore drifting to the northeast, and the coast curves out to a low limestone promontory at Cabo Salou. The coastal lowlands revive at Tarragona, along the Costa Daurada. There are sectors of low limestone cliff and gravelly beaches extending to the seaside town of

Sitges. This was one of the early resorts, established through the efforts of the Society to Attract Foreigners in the 1920s. The sandy beach has been renourished and maintained by many nearshore breakwaters (rompeolas), behind which are little tombolos, and between which are low, narrow and locally gravelly sectors (> Fig. 8.14.8). At the eastern end is the old town of Sitges behind a rocky promontory. The steep Costas de Garraf, east of Sitges, is on Cre­ taceous sandstones. It is trenched by short, deep valleys

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⊡⊡ Fig. 8.14.9 Sand is being extracted from the beach at Platya de Castell to prevent it building up and spilling into an adjacent marina. (Courtesy Geostudies.)

⊡⊡ Fig. 8.14.10 Sand-barred distributary of the Lobregat River, near Barcelona. (Courtesy Geostudies.)

descending to coves. At Castelldefels the steep slopes pass inland behind the broad plains of the Llobregat delta. Sand drifts eastward, and at La Pineda it has been intercepted by a breakwater to form the wide Platya de Castell beach, which is being denourished to reduce sand spill into a marina (> Fig. 8.14.9). To the north the sandy shore of the Llobregat delta has a wide beach (> Fig. 8.14.10) backed by ­pinewoods and suburbs on low dunes, extending to

the Barcelona waterfront. In the hinterland are segments of Pleistocene (Tyrrhenian) beaches, while the Llobregat delta has a sequence of Holocene deposits extending at least 54 m below present sea level. North-east of Barcelona the coast is generally steep, with beaches fringing a narrow coastal plain which widens at valley mouths, as at the Rio Tordero. Inland are the scarped ridges of the Barcelona lowland. Beyond Blanes is

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the embayed coast of the Costa Brava, with bold cliffed promontories of blocky granite between curving bay beaches, as in the Golfo de Rosas. These correspond with valleys drained by rivers such as the Duro, the Ter, the Fluvia and the Mugo, which have formed alluvial lowlands but not protruding deltas. To the north the high semi-arid promontory of Cadaques has small coves and steep-sided inlets (calas) with gravel and sand beaches between headlands of Mesozoic rock extending out to Cabo Creus. The orientation of the bays and headlands is related to jointing and faulting as well as lithological variations. The irregular coastline continues across the strongly-folded structures of the eastern Pyrenees to the French border at Port Bou.

5.  The Balearic Islands The Balearic archipelago represents an eastern continuation of the Betic Cordillera of southernmost Spain, with topography following a structural SW–NE orientation. Ibiza, the western island, is linked by a submarine shoal to Formentera to the south, and the main island of Majorca has similar undersea links to Cabrera, to the south, and Menorca, to the east. The northern coasts of Ibiza, Majorca and Menorca are steep and cliffy on Jurassic-Cretaceous limestones showing northward folding and thrusting. They

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have many steep-sided marine inlets (calas) and valleymouth bays between rocky headlands. There are trottoirs (visors of calcareous algae) above notches and narrow shore benches. Menorca is generally low-lying, and the northern cliffs cut across segments of the Palaeozoic basement, including Devonian schists. The Mesozoic formations are bordered southward in these islands by less strongly folded Tertiary and Quaternary sediment. On Majorca larger bays open to the southwest (Badia de Palma) and northeast (Badia de Pollensca, Badia d’Alcudia), fringed by gently curving sandy beaches. The southwest coast of Menorca is relatively straight and incised by numerous sub-parallel river valleys descending to calas. An artificial beach has been emplaced on the coast of Ibiza.

References Cotton CA (1956) Rias sensu stricto and sensu lato. Geogr J 122:360–364 Guilcher A (1974) Les rasas, un probleme de morphologie littorale generale. AnnGeogr 83:1–33 Marques MA, Julia R (1985) Spain. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, Stroudsburg, Pennsylania, pp 397–410 Trenhaile AS, Perez Alberti A, Martinez Cortizas A, Costa Casaia M, Blanco Chao R (1999) Rock coast inheritance: an example from Galicia, northwest Spain. Earth Surf Proc Land 24:605–621

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8.14.1 Gibraltar

Eric Bird

1.   Introduction Gibraltar is a high and narrow limestone peninsula 5 km long from north to south and up to 1.5 km wide, rising to a height of 421 m at Rock Gun. It originated as a double tombolo, attached to mainland Spain by a sandy isthmus containing a lagoon that has been reclaimed, and it has a coastline 12 km in length. The west coast, facing the Bay of Algeciras has a harbour sheltered by large breakwaters and an airport runway protruding into the sea; southwest is the Strait of Gibraltar, opening to the Atlantic Ocean. The westward slopes are steep and terraced (partly artificially) while the eastern side has great cliffs descending to the Mediterranean Sea (> Fig. 8.14.1.1). The limestone is of Lower Jurassic age, which moved down from the east over Upper Jurassic shale beds and Eocene sedimentary rocks in a tectonic displacement (allochthonic, or transported within a geosyncline) resulting in older formations resting upon younger. This inversion may have been the outcome of crushing between the African and Iberian tectonic plates (Bailey 1953). Within,

the limestone rock has been dissolved by underground streams to form caves, as at St. Michael’s, with prominent dripstone formations. On the crest of the range there are indications of frost shattering during Pleistocene cold phases. The climate is Mediterranean, with warm summers and mild winters. Mean annual temperatures range from 24°C in August to 12.8°C in January, and the annual rainfall averages 907 mm. The prevailing winds are westerly, but large waves are unusual in the Bay of Algeciras and Atlantic swell is much refracted before it reaches the west coast of Gibraltar. Occasional easterly winds produce waves from that direction, but there is little longshore drifting. Tides are diurnal, and the mean spring tide range 0.9 m.

2.  The Gibraltar Coastline A sandy beach runs along the western coast of the isthmus, much modified at the southern end by the protruding airport runway and by land reclamation. Gibraltar

⊡⊡ Fig. 8.14.1.1 The plunging cliffs of Gibraltar, seen from the south. (Courtesy Geostudies).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.14.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Harbour presents an artificial coastline, with reclamation in the dock area behind the breakwaters. To the south low cliffs appear, with basal notches and ramps. A mole shelters Rosia Bay, in which there is a small sandy beach. The cliffs increase southward round Camp Bay to Little Bay, both of which have sandy beaches. The southern end of the peninsula has cliffs backed by steep slopes which are fortified by walls. The cliffed coast curves round to the cove at Dead Man’s Beach and the headland at Europa Point, on which stands the lighthouse marking the eastern end of the Strait of Gibraltar. On the east coast the limestone cliffs rise sharply, following a major fault line, and the shore is rocky, with small notches and abrasion ramps and some karstic features produced by sea spray weathering. A corniche road follows the coast, tunnelling behind the great limestone wall, 300 m high, that truncates the southern end of the range, and to the north steep slopes descend to cliffs and rocky shores, with segments of sandy beach. The conspicuous Surface Water Catchment is a slope of corrugated iron that sheds rainwater into an underground system to

­ rovide Gibraltar, otherwise without springs or streams, p with a water supply. Below is Sandy Bay, and narrow beaches extend along North Sandy Bay and Blackstrap Cove. Shirley Cove is rocky, and Guilds Point, Blair’s Point and St Abbe’s Head are minor protrusions. The eastern slopes have remnant aprons of aeolian sand, the origin of which is uncertain: it may have been blown up from wider beaches, perhaps when sea level was lower, or it may have come across from the Sahara. Catalan Bay has a recreational beach with a promenade, a lido and a village, and above it the cliffy limestone range comes to an end overlooking the sandy isthmus towards Spain. North of Catalan Bay is long Eastern Beach, running along the eastern shore of the sandy isthmus into Spain and the Costa del Sol.

Reference Bailey EB (1953) Notes on Gibraltar and the Northern Rif. Q J Geol Soc Lond 108:157–176

8.15 Portugal

Carlos Morais*

1.  Introduction The Portuguese coastline is about 845 km long. The climate is cool and maritime, with a transition from cool temperate in the north to Mediterranean in the south. The summer is dry, the winter wet, especially in the north. Porto has a mean monthly temperature of 9°C in January and 19°C in July and a mean annual rainfall of 1,151 mm; Lisbon has a mean monthly temperature of 11°C in January and 22°C in July and a mean annual rainfall of 686 mm. Tides are semidiurnal, and mean spring tide ranges typically mesotidal: Barra de Aveiro has 2.5 m, Figueira do Foz 3.0 m, Peniche 3.0 m, Lisbon 3.3 m, Setubal 4.0 m, and Cabo San Maria 2.6 m. Currents in coastal waters flow from north to south. The west-facing coast is exposed to Atlantic swell and waves generated by westerly and northwesterly winds in coastal waters; in winter, the southwesterly component is also strong, especially north of Cape Carvoeiro. The southern (Algarve) coastline, east from Cape St. Vincent, receives southerly ocean swell and local wind-generated waves from the southwest (especially in spring and early summer) and southeast (in late summer and autumn). There are higher mean annual wind velocities (about 16 km/h) on the western coast than in the Algarve (6–9 km/h). In general, wave energy decreases from north to south (Barceló 1975). In response to these processes, predominant longshore drifting is generally from north to south along the western coastline and from west to east in the Algarve. Reversals occur in the lee of major promontories south of Lisbon and in the Bay of Setubal, where shelter from northwesterly waves permits a local dominance of northward drifting. Castanho et al. (1981) have estimated net annual littoral drift at 150,000 m3 at Leixoes, 1 million m3 at Aveiro, 200,000 m3 south of Figueira da Foz, and 40,000 m3 at Quarteira.

2.  The Portuguese Coastline Features of coastal geomorphology from north (Minho River) to southeast (Guadiana River) are described in sequence. The estuary of the Minho River on the Galician (Spanish) border is long and narrow, in contrast with the

wide and deep rias further north. There is a submarine canyon offshore, and it is possible that the continental shelf has here escaped the tectonic subsidence that has taken place off Galicia. The coast south to Porto (Oporto) is low-lying, with sandy beaches and rocky sectors with nearshore reefs between the mouths of small rivers. The narrow coastal plain is backed by steep granitic slopes of an ancient massif. The Lima River at Viana de Castelo also has a narrow estuary and a submarine canyon offshore. At Esposende, grassy dunes back a sandy beach, Praia de Suave Mar, north of the Rio Cavado. Predominant southward drifting of beach sediment has resulted in interception to prograde the shore north of the Leixoes oil terminal at Porto, and as a result there has been erosion of beaches to the south at Espinho. The mouth of the River Douro is partly enclosed by a spit, impeding access to the harbour. To the south, the coast is low and sandy, with coalescent spits forming barriers, capped by coastal dunes. The barriers are interrupted by an artificial entrance, bordered by protruding breakwaters, to the Ria de Aveiro, a 60 sq. km lagoon consisting of the converging estuaries of several rivers with an artificial entrance giving access to Aveiro Harbour. Sand drifting southward has accumulated as a wide beach north of the breakwaters, while to the south the beach at Praia de Barra is narrow and the backshore dunes eroded. Tidal flats are extensive here, rising to salt marshes and low marshy islands. The Ria da Costa Nova is a long, narrow lagoon behind the coastal barrier: it led to variable outlets to the south before the artificial entrance was cut. To the south, the dune fringe has been cut back, as at Praia de Vagueira, and there are several seaside resorts, some with concrete esplanades as at Praia de Mira. Beyond the end of the Ria da Costa Nova, the dune fringe widens, and the dune vegetation is sparse, as at Praia de Tocha (>Fig. 8.15.1). The sandy beach comes to an end at Cape Mondego. Cape Mondego is a monoclinal Jurassic formation with a basal cliff in stratified grey mudstones and upper slopes mantled with angular gravel in an earthy matrix. This is a Pleistocene periglacial deposit similar to those seen on slope-over-wall coasts along the Atlantic seaboard in Ireland, Wales, Cornwall, Brittany, and Galicia, and

*Edited version of a chapter by Carlos Morais in The World’s Coastline (1985: 411–417). All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.15, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.15.1 Dune barrier, Praia de Tocha. (Courtesy Geostudies.)

appears to be the southernmost limit of this feature. Ribs of mudstone run out across the shore, separating sandy coves, and there is a large coastal quarry, from which rubble has spilled down on to the shore. The headland shelters Buarcos Bay and the harbour of Figueira da Foz from northwesterly storms. Beyond Buarcos, the beach widens, and at Figuera da Foz sand has accumulated alongside breakwaters at the mouth of Rio Mondego to form a beach more than 100 m wide. The Mondego estuary has rushy salt marshes and salt-making ponds. To the south, the coast fringes a sandy plain with extensive dunes, interrupted by a limestone promontory at Pedrogão and the stabilised outlet with the Rio Liz. The beach has lobate segments behind shore ­platforms and reefs, as at Pinhal de Leiria, and the coast rises to a headland on the high rocky massif at Sitio. The submarine canyon of Nazare heads a short distance ­offshore, with depths of 200 m close to the coast in front of Sitio. South from the high cliff at Sitio is the wide Nazare beach, and breakwaters enclose a harbour. Concha de Sao Martinho do Porto is an oval-shaped tidal lagoon behind a gap in a coastal ridge, through which waves are diffracted. Dunes on the inner shore have been stabilised by black hessian sheets. The coast steepens to Ponte dos Covinhos, where the coastal slope is dissected by gullies and chines. At the mouth of the Rio Obidos, another lagoon has an entrance frequently deflected by southward spit growth. There are shoals of inwashed rippled sand between ebb and flood channels, and bordering salt marsh.

The beach then curves out toward the Peniche peninsula. Baleal is a village on an island attached to the coast by a sandy tombolo and bordered by Ilha das Pombas, where the dipping limestones have been planed off (>Fig. 8.15.2), presumably by the sea when it stood at a higher level. Cape Carvoeiro to the west is a larger tombolo, with a dune-fringed sandy isthmus and the fishing harbour of Peniche enclosed by curving breakwaters on its more sheltered southern side. On the northern coast cliffs are cut across dipping strata (>Fig.  8.15.3), fronted by a broad intertidal platform. The western headland has bold cliffs in limestone showing karstic weathering, with tors and klints (>Fig. 8.15.4), and offshore are the flat-topped rocky islets of Berlengas. Peniche faces south, and the high cliffs continue, with relatively stable pocket beaches, river mouths as at Rio Sizandro, and small fishing harbours sheltered by break­ waters as at Ericeira. There is an escarpment cliff in landward-dipping strata at Foz do Lisandro (>Fig.  8.15.5), and the harder strata form structural ledges on cliffs as at Azenhas do Mar and from Praia das Matas to Cabo da Roca. The high cliffs at Cabo da Roca (144 m), the westernmost point of Europe, are on a granite batholith. They are bordered by hanging valleys and boulder-strewn rocky shores. The coast becomes lower to the south, with ­re-entrants cut out by marine erosion along basic dike outcrops. Praia das Macas is a sandy surf beach between rocky shores dominated by undulating limestone strata, which form shore ledges and scarps. The Sierra de Sintra rise inland, and there are shore platforms and pocket

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⊡⊡ Fig. 8.15.2 Planation surface cut across dipping limestone at Baleal, near Peniche. (Courtesy Philippa Barr.)

⊡⊡ Fig. 8.15.3 Cliffs north of Peniche. (Courtesy Geostudies.)

beaches extending past Cape Raso along the south-facing sector toward Lisbon. The Tagus estuary has a narrow mouth with the shoal of Bugio in the channel. It widens inward to the Mar da Palha, with the city and port of Lisbon on the north shore and shipyards at Lisnave to the south; it is essentially

lagoonal, with a delta and salt marshes at the inflow of the Tagus River. South of the Tagus estuary, beach erosion is a problem on the Caparica coast, resulting in the building of numerous groynes. A long, gently curving coastline, determined by the refraction of northwesterly waves past Cape Raso,

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⊡⊡ Fig. 8.15.4 Karstic weathering on cliff at Peniche. (Courtesy Geostudies.)

⊡⊡ Fig. 8.15.5 Escarpment cliff north of Foz do Lisandro. (Courtesy Geostudies.)

has a wide sandy beach with some dunes, backed by degraded cliffs and interrupted by the Albufeira lagoon. Cape Espichel is the western limit of the Serra da Arrabida, the southern flanks of which end in high cliffs. There are small sandy beaches, difficult of access, and ­fishing harbours such as Sesimbra. The large estuarine lagoon of the Sado has a marine outlet constricted by the northward growth of the Troia sand spit, and is bordered

by extensive salt marshes, dominated by Spartina grass. The port of Setubal stands on its northern shore. A long, gently curving sandy coastline extends south to the Cape of Sines, its outline determined by the ­refraction of northwesterly waves around Cape Espichel. Behind the beach are the Melides and Santo Andre lagoons, both with intermittent outlets to the sea through sandy dune-capped barriers. The Cape of Sines is a ­volcanic

Portugal

outcrop, beside which a harbour has been built. Southward to Cape St. Vincent, the coastline is generally cliffed, with rocky shore platforms and intermittent beaches, and some small estuaries, for example the Mila at Vila Nova de Milfontes. Cape St. Vincent and the nearby Ponta de Sagres are bold headlands with vertical cliffs on Jurassic limestone outcrops. The south-facing coastline of the Algarve has high cliffs cut in Miocene limestone with some karstic ­features and nearshore rocky reefs. It is interrupted by bay  beaches and by small spit-constricted river mouths, for example the Arade estuary at Pertimão. Eastward drifting of beach sand is evident at the artificial harbour of Vilamoura, where groynes have been inserted to retain beach material on the eroding coastline. At Praia da Rocha, an artificial beach has been placed in front of cliffs where no beach existed previously (Psuty and Moreira 1990).

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East of Quarteira, the coast is low and sandy, with dune-capped barriers enclosing lagoons and salt marshes, notably behind the cuspate foreland of Santa Maria, where the harbours of Faro and Olhao are situated. The beaches continue to the Spanish border at the mouth of the Guadiana River, the sector east of Faro being subject to occasional levante winds blowing from Gibraltar.

References Barceló JP (1975) On the Portuguese wave regime. Pro 14th Coast Eng Conf 112–131 Castanho JP, Gomes NA, Mota Oliveira IB, Simoes JP (1981) Coastal erosion caused by harbour works on the Portuguese coast and corrective measures. Proc of the 25th P I A N C, Section 2(15):877–898 Psuty NP, Moreira ME (1990) Nourishment of a cliffed coastline, Praia da Rocha, the Algarve, Portugal. J Coastal Res Special Issue 6:21–32

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8.16 Italy

Paolo Ciavola

1.  Introduction The Italian coastline is about 7,600 km long, including Sardinia and Sicily. Much of Italy has a warm temperate (Mediterranean) climate, with mild rainy winters and warm dry summers, but the Adriatic coast is cooler and drier. On the west coast, Florence has mean monthly temperatures of 5.6°C in January rising to 25° in July and an average annual rainfall of 901 mm, while Rome has 7°C in January and 25°C in July, with an average annual rainfall of 657 mm. On the east coast, Venice has 3.3°C in January and 23.9°C in July, and an average annual rainfall of 725 mm. Westerly winds predominate, but there are occasional hot southerlies in the south (scirocco) and cold north-easterly (bora) winds in the Adriatic Sea. Waves approach the coast from these various directions, according to aspect and configuration. Tides are very small, up to 0.2 m around much of the coast but larger in the northern Adriatic, where Venice has 0.8 m. There are larger fluctuations of sea level in relation to onshore and offshore winds and variations in atmospheric pressure, with occasional storm surges, particularly in the northern Adriatic where Venice is subject to increasing sea flooding by the acqua alta (Ciavola et al. 2002). The mountain ranges produced by Alpine folding that run through Italy produce steep coasts in Liguria, Calabria, and Sicily, with spurs that form hilly promontories, and many intervening coastal lowlands, the most extensive being in the northeast, north, and south of the Po delta. Holocene marine submergence has affected the coasts of Sardinia, Liguria, and Carso near Trieste, but generally drowned valley mouths have been infilled with sediment. Pleistocene and Holocene tectonic uplift has occurred on parts of the Tyrrhenian, Ionian, and Adriatic coastlines, producing marine terraces at various levels, but these have also been influenced by eustatic oscillations of sea level, and it is difficult to distinguish these effects. In general, uplift was later along the Adriatic and Ionian coasts than along the Tyrrhenian. It is currently believed that the highest and oldest marine terraces were formed in the Lower Pleistocene. Along the margins of coastal plains (that of the Po valley in particular), tectonic subsidence is evident, and there are submerged Pleistocene and Holocene

f­ eatures on the bordering sea floors, including coastlines, beaches, dunes, and river beds. Examples may be found in the gulfs of southern Sardinia, La Spezia, Manfredonia, Venice, and the Pontine littoral zone. There are extensive steep rocky coasts, some with stable bluffs, others receding cliffs, where the zone between the 100 m land contour and 100 m sea floor contour is narrow, and intervening low-lying coastal plains where it is wide. An abundant fluvial sediment yield has produced several substantial deltas, notably those of the Tiber and the Po. Beaches are very extensive, except on a few sectors of plunging cliffs. Italian beaches have received much of their sandy and gravelly sediment from rivers, but some have been derived from cliffs, or from the sea floor. Comparisons with historical maps have shown the pattern of erosion and accretion over a century (from 1863–1892 to 1953–1972), and during the past few years almost all Italian beaches have receded, largely because of human activities such as the building of dams on sediment-yielding rivers, the quarrying of sand from the shore, the construction of breakwaters, and coastal land subsidence induced by groundwater extraction (Zunica 1976, 1985). It is evident that about half the coastline is receding, often at dramatic speed. Beach replenishment has only started relatively recently in the country, with major replenishment schemes undertaken initially along the Venice littoral and in more recent times along the coast of Lazio and Emilia-Romagna.

2.  The Italian Coastline From the French border near Ventimiglia, the Ligurian coast is steep, close to Appennine watersheds, with short rivers descending incised valleys to small deltaic lowlands. Fluvial sediment yields have increased with deforestation of the catchments, augmenting the supply of sand and gravel to beaches and causing delta growth, as in the Argentine delta (Fabbri and Bird 1993). There are variable patterns of longshore drifting at river mouths, but the sea floor slopes away steeply, and much sediment is lost offshore because of the presence of deep canyons related to river mouths (Ferretti et al. 1992). Beaches are

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⊡⊡ Fig. 8.16.1 The beach at Alassio on the Ligurian coast. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.2 Plunging cliffs in steeply-dipping rocky formations near Vernazza on the Ligurian coast. Notice the debris fans at the base of slopes.

generally narrow on western side (>Fig.  8.16.1), and often absent in eastern sector at the base of plunging cliffy headlands (>Fig. 8.16.2) or present in the form of small pocket beaches, mainly composed of coarse sediment. Exposed to storm waves from the south, they are subject to intense erosion. There has been reclamation near coastal cities such as Genoa.

There are deltaic lowlands at La Spezia, and on a larger scale inside the Tuscany coast south of Viareggio, where the Arno has built a wide plain at Pisa (Mazzanti and Pasquinucci 1983). Coarse pocket beaches are present along the coastline of the Elba Island, while the other major island of the Arcipelago Toscano, the Giglio Island, is dominantly shaped by granite cliffs (>Fig. 8.16.3), with

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⊡⊡ Fig. 8.16.3 Granite cliffs on Giglio Island.

⊡⊡ Fig. 8.16.4 The double tombolo at Orbetello. (Courtesy Geostudies.)

the only beach located at Campese. Sandy beaches fringe the lowlands along the coast of the mainland, but a decreased sediment supply by rivers has generated coastal erosion that prompted the building of numerous nearshore breakwaters, which divide the beach into a series of ­compartments between cuspate spits. Longshore drifting has deflected the mouths of Tuscan rivers (Pranzini 2001). To the south, erosion has been extensive along the

Latium coast (Caputo et al. 1983). A hilly promontory at  Piombino juts out toward the high island of Elba, and  south of the Ombrone delta, Monte Argentario (635 m) is attached to the mainland by a double tombolo at Orbetello (Tombolo della Gianella to the north, Tombolo di Feniglia to the south, enclosing a lagoon (>Fig. 8.16.4). Broad bays are backed by alluvial ­lowlands to Civitavecchia, and the Tiber downstream from Rome

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⊡⊡ Fig. 8.16.5 The ruins of Nero’s Villa fronting the cliffs near Nettuno. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.6 Beach accretion beside the harbour breakwater at Terracina. (Courtesy Geostudies.)

reaches the sea on a broad deltaic plain that has fringing beaches, beach ridges, and dunes. At Anzio, there are cliffs in Tertiary sedimentary formations. At one point, the ruins of Nero’s villa, built against the cliff 2,000 years ago when sea level must have been lower, are being dissected by wave erosion (Bird and Fabbri 1987) (> Fig. 8.16.5). Torre Asturia is a rocky promontory separating long sandy beaches. The coastal plain of Latina and Sabaudia is notable for a dune-capped coastal barrier and

lagoons, with landward remnants of parallel Pleistocene barriers and dunes, which are unusual in Europe. The dunes are well vegetated, some with pine forests. The barrier comes to an end at the calcareous cliffy promontory of Monte Circeo. The Ponziane islands offshore have steep rocky coasts with some eroding cliffs and only minor beaches in coves, the largest Chiaia di Luna on the western side. There are high crumbling cliffs of grey and yellow weathered sandstone on the Isla Palmarola, where there is

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⊡⊡ Fig. 8.16.7 A zone of marine organisms on the columns of the Temple of Serapis indicates a phase when sea level was higher, relative to the land. (Courtesy Geostudies.)

underwater evidence of neotectonic movements. East of Monte Circeo, a sandy beach extends to Terracina, where there has been local accretion of westward drifting sand beside harbour breakwaters (>Fig. 8.16.6). The hilly spurs of Monti Aurunci separate the alluvial lowlands of Fondi, and a steep coast extends from Sperlonga to the tombolo at Gaeta. The hilly coast continues behind the Golfo di Gaeta, curving past Formia to Scauri, where it gives place to a broad alluvial lowland. Here, the Garigliano River crosses a deltaic plain, and the spur of Mount Massica is followed south of Mondragone by the little Volturno delta. The lowland comes to an end at the Monte di Procida promontory, and offshore is the high island of Ischia, which shows evidence of Holocene uplift. The tufaceous Procida promontory and island mark the beginning of the Bay of Naples, and parts of the Neapolitan coast show evidence of land movements, both upward and downward, sometimes rapid and of considerable extent, connected with the volcanic nature of this area. The Temple of Serapis at Pozzuoli (>Fig.  8.16.7) is notable for evidence of a rise and fall in sea level, due to tectonic movements associated with nearby Vesuvius, as noted by Lyell in 1830 and Huxley in 1877. The columns show a horizon with marine borings, which formed while the temple was partly submerged between the fifth and ­fifteenth centuries ad and emerged when it was subse­ quently uplifted. The steep coast of Pozzuoli extends eastward to Naples, which stands at the northern end of the bay. It has an

esplanade with a sea wall protected by a 10 m zone of limestone blocks, and no beach, except for a short sector at Rotunda Diaz where a small nearshore breakwater protects a sandy tombolo. The beaches of the Bay of Naples are generally of dark grey volcanic sand, derived from the Vesuvius volcano, which rises steeply to 1277 m in the hinterland. To the south, the mountainous limestone peninsula of Sorrento extends out toward the Isle of Capri. A steep coast borders this upland, dissected by deep narrow valleys such as that of the Picentino River and extending round to Salerno. The Gulf of Salerno is backed by an alluvial plain extending south from Battipaglia on either side of the Sele River. There is a long gently curving sandy beach, backed by low dunes, with a small sandy cuspate delta at the mouth of the Sele. The coastline has been shaped by southwesterly waves arriving from a long fetch south of Sardinia and across the Tyrrhenian Sea. The alluvial plain narrows southward to Agropoli, and a steep, high promontory then runs out to Punta Licosa. On its southern side, a wide beach at Asceo is backed by a dune ridge. The Campania peninsula rises to Monte Bulgheria, and is incised by the Alento valley. The river mouth is deflected southward, and there are sandy ridges on its southern side. The steep coast continues round to Policastro. There are several small seaside resorts, with problems of depleted beaches. Sapri stands behind an inset bay, and to the south the steep, straight coast borders the Catena Costiera (coast range). The railway runs close to the coastline. Many short, steep

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streams descend deeply incised valleys from a divide a short distance inland. East of this, longer rivers flow down to the Crati delta on the east coast. The steep coast has an uplifted series of marine terraces. Rivers supply gravel to beaches, which is sorted and reduced in calibre as it drifts along the shore (Bartholoma et al. 1998). Beach erosion has become prevalent along this coast. There is a discontinuous, narrow, coastal lowland that fans out only near the main rivers (locally called “fiumare”), characterised by occasional damaging seasonal floods, typical of a torrential regime with consistent peak water discharges. High limestone cliffs around Maratea descend to a karstic shore. Isola di Dino is a 65 m high island just offshore to the south of Praia Mare. Ridges end in rocky promontories, as at Capo Scalea. Longshore drifting is southward. Many sectors have only narrow (if any) beaches, and numerous nearshore breakwaters have been built to protect them. At Diamante, the Corvino River has built a small delta, and south of Scalea the Lao and Abetamarco Rivers have blunt deltas. Steep slopes descend to the seaside resort of Paola, where eroded shores have been armoured with boulders. Near Amantea, there are occasional rocky stacks offshore. To the south, the River Savuto has a small sandy delta, and Capo Suvero is a low promontory. NW of Lamezia a sandy barrier beach encloses a narrow lagoon, Lago La Vota. To the south, coastline recession has left several buildings on protected promontories

between scoured bays. Beyond the deltaic plain of S. Eufemia, a steep coast runs from Pizzo out to Capo Vaticano and round to Nicotera. Rosarno stands on the Mesima alluvial plain, and Gioia Tauro and Palmi are small resorts with beaches protected by nearshore breakwaters. The coast becomes steep and mountainous past Bagnara Calabra, with rocky cliffs on headlands and beaches in valley-mouth bays. It continues beside the Strait of Messina (with Sicily to the west), and round to the southern coast of Calabria, facing the Ionian Sea. The southern and western coasts of the Gulf of Taranto are bordered by steep slopes descending to a narrow beach-fringed coastal plain. Because of the fluvial supply, beaches are often composed by gravel and coarse sand. There are limestone areas with partly drowned dolines. The steep slopes show up to seven marine terraces, the highest 400 m above sea level. Several rivers, including the Sinni, Agri, and Cavone, flow across a narrow coastal plain to small lobate deltas on the north-west coast of the Gulf, and the east coast, beyond Taranto, is a generally low-lying hilly peninsula culminating in the Capo San Maria de Leuca. The whole of the Taranto area has a welldocumented history of late Pleistocene coastal changes and tectonism (Belluomini et al. 2002). On the Adriatic coast from Brindisi to the Monte Gargano, promontory high rocky coasts prevail, with cliffs cut in biogenic calcarenites, with basal notches and few ⊡⊡ Fig. 8.16.8 High cliffs cut in marls and limestones on the Conero Promontory.

Italy

beaches. However, the Gulf of Manfredonia is backed by a coastal plain with wide beaches receiving sediment from the Ofanto and smaller rivers. The Gargano promontory shows evidence of Holocene uplift and historical sea level changes (Mastronuzzi and Sanso 2002). There are flat-topped caves cut into horizontally stratified Jurassic limestone. North of this promontory, sand and gravel barriers enclose the coastal lagoons of Varano and Lesina and continue as beaches interrupted only by small stretches of high steep coast (Parea 1983). The beaches are fed with sand and gravel by numerous short rivers descending from the Appennine ridge. They have incised valleys into clays, gravels, and sands, and the coastal slopes show up to five series marine terraces. This type of coast extends north to the Conero headland, formed by marls and limestones with slope instability problems (>Fig.  8.16.8). The cliffs form sea stacks and small coastal resorts are located at their base. Northward of Ancona, the coastline becomes sandy again, but northward of Pesaro a new raise in landscape is present, with the San Bartolo area again occupied by cliffs with coarse beaches at their base (>Fig. 8.16.9). At Pesaro, the Appennine slopes pass inland south of the broad Plain of the Po River (Pianura Padana). The coast past Rimini to Ravenna is low-lying, fringed by beaches and dunes. Dune ridges have disappeared during post World War II economic development of the coastal zone. One of the few areas where dunes survive is between Lido di Classe and Lido di Dante (>Fig. 8.16.10). Beach erosion is partly the outcome of coastal land subsidence following extraction of groundwater, oil, and natural gas in the Ravenna area (Gambolati et al. 1999). It has been countered by building numerous groynes (>Fig. 8.16.11) and nearshore breakwaters (>Fig. 8.16.12), some of which have induced lee-side cusps, which may grow into tombolos dividing the beach into rounded bays. The beach becomes a barrier enclosing coastal lagoons at Comacchio. To the north is the large Po delta, with distributary channels that have varied in size and location historically since pre-Etruscan times. At one stage, the river was diverted from a northward outlet toward the Venice Lagoon (Fabbri 1985). The Po drains a basin of more than 70,000 sq. km and its delta extends out into the sea for about 20 km. On the delta, shore sandy beaches are interrupted by sectors of reedswamp around river mouths, where large spits protrude into the Adriatic Sea (>Fig. 8.16.13), often enclosing coastal lagoons (sacche) like at Goro and at Canarin. In these lagoons, mollusc farming has become a regular practice, providing an important revenue for the local economy.

8.16

⊡⊡ Fig. 8.16.9 Steep coast and shingle beach south of Vallugola. (Courtesy Geostudies.)

Beaches continue northward, becoming barrier islands in front of the Lagoon of Venice, with some low dunes. The lagoon is bordered by extensive salt marshes (barene) (>Fig. 8.16.14), large portions of which have been enclosed and reclaimed, but erosion has been extensive in recent decades (Cavazzoni 1983). Water exchange inside the lagoon is ensured by the presence of three inlets (Lido, Malamocco, and Chioggia). The inlet of Lido is connected to a deep navigation channel that serves the oil and chemical industry at Marghera, where the ship traffic generates consistent bed resuspension (Ciavola 2005). Segments of the lagoon have been enclosed to form fishponds (valli). The coastal plain consists of alluvium deposited by the Adige, Brenta, Piave, Livenza, Tagliamento, and Isonzo Rivers. The Venetian beaches have been much influenced by human activities. From the thirteenth century onward, many river courses (including that of the Piave) were

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⊡⊡ Fig. 8.16.10 Dune field between Lido di Classe and Lido di Dante. Notice in the background one of the numerous oil platforms drilling for gas.

⊡⊡ Fig. 8.16.11 Groynes at Rimini have failed to maintain the beach. (Courtesy Geostudies.)

diverted to reduce siltation in the Lagoon of Venice and thus maintain it as a defensive moat for the city. Erosion of  the barrier islands culminated in the construction of murazzi (sea walls built during the eighteenth century by the Venetians), then long, protruding stone jetties to stabilise the navigable entrances. These interrupted longshore

drifting and resulted in further erosion. As a result, parts of the coastline are now heavily armoured by sea walls (>Fig. 8.16.15). Nowadays, the beaches along the barrierislands are replenished using offshore sands. Tides are of considerable importance along this part of the coast. In the Adriatic, which is practically a closed

Italy

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⊡⊡ Fig. 8.16.12 Nearshore breakwaters on the Rimini coast. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.13 The southern coast of the Po delta, with barrier islands and spits formed at the mouth of a major distributary. (Courtesy Geostudies.)

basin, they may reach 1 m. Because of the effects of the scirocco winds (from the southeast), low atmospheric pressures, and seiches, sea level may also rise by almost 2 m. The impacts of such storm surges have become more frequent in recent years, largely as the result of land sub­ sidence due to tectonic depression and groundwater extraction, and have caused flooding by exceptionally high tides, known in Venice as acqua alta (Pirazzoli 1983). These storm surges may cause serious damage to the barrier

islands and the Venetian hinterland, partly because the defences are not of sufficient dimensions to deal with these anomalous conditions. Mean tide range in the lagoon is about 0.80 m, and when the water level rises 0.80 m above mean sea level it floods St. Marks Square. In 2000, it reached at least 1.00 m on ten occasions. Sirens warn of the major flooding that occurs when the level exceeds 1.10 m, now about four times a year, and there is a risk that the devastating floods of 4 November 1966, which attained 1.94 m, will

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⊡⊡ Fig. 8.16.14 Eroding salt marshes in the Venice Lagoon. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.15 After a sea wall and groynes failed to retain the beach at Litorale di Pellestrina on the Venice coast, limestone blocks were dumped to form an artificial coastline. (Courtesy Geostudies.)

recur. Currently, flood gates are being built at the inlets (the Moses Project) inserting hydraulic structures in the three entrances that can be raised to exclude the sea during storm surge episodes. It is not possible to close the entrances because ships require access to Porto Marghera, and because there are also problems of lagoon pollution, especially from the industries of Porto Marghera.

There is a long gently curved beach-fringed coastline between Venice and Trieste. The lagoon of Grado and Marano is the second largest coastal lagoon in the Adriatic and in many ways resembles the larger Venice Lagoon. How­ever, unlike in Venice where the continuous inlet dreging has created anomalous deep tidal inlets, here the mouth of the lagoon has large ebb tidal deltas. Just before

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Trieste the beach ends, and cliffs and bluffs border an indented coastline.

References Bartholoma A, Ibbeken H, Schleyer R (1998) Modification of gravel ­during longshore transsport, Bianco Beach, Calabria, Southern Italy. J Sediment Res 68:138–147 Belluomini G, Caldara M, Casini C, Cerasoli M, Manfra L, Mastronuzzi G, Palmentola G, Sansò P, Tuccimei P, Vesica PL (2002) Age of  late  Pleistocene shoreline, morphological evolution and tectonic ­history of Taranto area, Southern Italy. Quat Sci Rev 21(4–6):425–454 Bird ECF, Fabbri P (1987) Archaeological evidence of coastline changes illustrated with reference to Latium, Italy. In: Trousset P (ed) Déplace­ ments des lignes de rivage en Mediterannée, Colloques Internationaux, C.N.R.S., pp 107–113 Caputo C, D’Alessandro L, La Monica GB, Landini B, Lupia Palmieri E, Pugliese F (1983) Erosion problems on the coast of Latium, Italy. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean Sea, Bologna, p 59–68 Cavazzoni S (1983) Recent erosive problems in the Venetian Lagoon. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean Sea, Bologna, p 19–22 Ciavola P, Organo C, Leon Vintró L, Mitchell PI (2002) Sedimentation processes on intertidal areas of the Lagoon of Venice: identification of exceptional flood events (acqua alta) using radionuclides. J Coastal Res Special Issue 36:139–147

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Ciavola P (2005) Sediment resuspension in the Lagoon of Venice: ­short-term observations of natural and anthropogenic processes. Z Geomorphol Special Issue 141:1–15 Fabbri P (1985) Coastline variations on the Po delta since 2,500 B.P. Z Geomorphol, Supplementband 57:155–167 Fabbri P, Bird ECF (1993) Geomorphological and historical changes on the Argentine delta, Ligurian coast, Italy. GeoJournal 29:428–439 Ferretti O, Immordino F, Niccolai I (1992) Transport and distribution of sediments along the Ligurian coast. Hydrobiologia 235–236:17–32 Gambolati G, Teatini P, Tomasi L (1999) Coastline regression of the Romagna region, Italy, due to natural and anthropogenic land subsidence and sea level rise. Water Resou Res 35:163–184 Mastronuzzi G, Sanso P (2002) Holocene uplift rates and historical rapid sea level changes at the Gargano Promontory, Italy. J Quat Sci 17:593–606 Mazzanti R, Pasquinucci M (1983) The evolution of the Luni-Pisa ­coastline. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean Sea, Bologna, p 47–58 Parea GC (1983) The evolution of the Adriatic coast between the Tronto River and Rodi Garganico, Italy. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean sea, Bologna, p 39–46 Pirazzoli PA (1983) Flooding (acqua alta) in Venice (Italy): a worsening phenomenon. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean Sea, Bologna, p 33–38 Pranzini E (2001) Updrift river mouth migration on cuspate deltas: two  examples from the coast of Tuscany, Italy. Geomorphology 38:125–132 Zunica M (1976) Coastal changes in Italy during the past century. Italian Contrib to the 23rd Int Geogr Congr 275–281 Zunica M (1985) Italy. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 419–429

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8.16.1 Sardinia

Eric Bird

1.  Introduction

2.  The South and East Coast

The island of Sardinia (Sardegna) includes a Hercynian (Variscan) massif of mainly crystalline composition, with associated Mesozoic and Tertiary sedimentary rocks, limestones, and volcanics (Cocozza and Jacobacci 1975). Like Corsica, it has been separated from similar structures in southern France on tectonic microplates that have rotated as they separated from the Riviera region. The various rock formations trend mainly N–S and NW–SE. Lithology and structure have influenced the coast, which has sectors of high, rocky cliff on a variety of igneous, metamorphic, and sedimentary formations. Pleistocene dune calcarenite is present on some cliffy coasts, but shore platforms are poorly developed. There are several very high cliffs where Jurassic and Cretaceous limestone massifs plunge into deep nearshore water, and although some rock falls occur there has not been sufficient wave energy to have shaped these huge features. They are the outcome either of faulting or of partial submergence of mountainous areas. Parts of the east coast are also steep, with a sea floor declining abruptly offshore, and the dissected granite batholith that forms the northern peninsula has a ria coast. There is a lowland behind the Gulf of Cagliari in the south, extending as the Campidano rift valley floored with Quaternary sediment NNW to the Gulf of Oristano, and in the south-west is an upland of mainly Palaeozoic rocks. Curving sandy beaches border the larger coastal plains, and on parts of the coast become dune-capped barriers fronting coastal lagoons (Guilcher 1983). The enclosing barriers are low and narrow, and large waves that often overtop them have formed numerous washover fans. Seagrasses (Posidonia) has been washed up on the shore, and frequently rolled into balls by wave action. Tide ranges are small around Sardinia (0.2 m at Cagliari) and are sometimes exceeded by sea level fluctuations related to changes in barometric pressure or the effects of onshore and offshore winds.

Cagliari stands behind the peninsula of Capo Sant Elia, and from Poetta a curving 8 km long white sandy beach fronts a barrier enclosing lagoons and marshes, Spagno di Quartu. There are segments of a Pleistocene inner barrier at Is Arenas (Ulzega and Ozer 1980). To the southeast hilly granitic country runs out past Villasimius (where a weathered rock is in the shape of a shark’s fin) to Capo Carbonara, where the coastline turns northward. Offshore are Isola dei Cavoli and Isola Serpentara, with rocky shores and minor beaches. The hilly coast continues north to the Costa Rei, a lowland with residual lagoons, fringed by a curving sandy beach. Cape Ferrato is a granite headland, and to the north another curving sandy beach barrier runs to the mouth of the Flumendosa River, and on to Porto Corallo. The coast is then cut into hilly country on Ordovician and Silurian schists north to Capo San Lorenzo and the bay to Torre di Murtas. Granite outcrops between Torre di Murtas and Capo Palmeri, then the schists return, and there is more granite between Capo Sferracavallo and Sa Perda Pera. Another curved beach backs an embayment north to Punta su Mastixi, which is on andesite lava, and there are alternations of red porphyritic rock, dark diorite, and granitic veins in headlands at Capo Bellavista, near Arbatax, with intervening sandy bays. A coastal lowland with a sandy barrier enclosing the lagoon Stagno Tortoli extends to S. Maria Navarrese, followed by cliffs and bluffs in Palaeozoic schists and Siluro-Devonian calcareous sedimentary rocks, as at Capo di Monte Santu. North of Punta Pedro Langa, a short sector of granite coast is backed and bordered by stratified Jurassic rocks, which pass beneath massive limestone toward Capo di Monte Santo. This marks the beginning of the wide Golfo di Orosei, backed by steep, generally vertical cliffs of seaward-dipping Jurassic limestone extending north past Cala di Luna to Caletta di Fuili. Cala di Luna has a beach fronting cliffs that have several caves, and to the north the Grotta del Bue Marino (Cave of the Sea Oxen, i.e. rare seals) is cut

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.16.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8.16.1

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⊡⊡ Fig. 8.16.1.1 Vertical cliffs cut in piedmont gravels. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.1.2 Cala Gonone. (Courtesy Geostudies.)

into the cliff base. At Caletta di Fuili, a deep gorge with a gravelly watercourse opens to a beach of pale shingle. Offshore, a submarine canyon runs out to the east. Near Calette di Fuili, the limestone cliff has a basal skirt with close vertical rilling, undercut by a notch (>Fig. 8.16.1.1). The adjacent beach has limestone and basalt boulders, and well-rounded but poorly sorted shingle. A classic notch and visor profile is seen at the base of the limestone cliff.

The cliffs then pass northward behind a lower sector where they are fronted by Pleistocene olivine basalt and Holocene gravel piedmont fans of periglacifluvial origin. In a central sector, the piedmont gravels are incised by a gorge that opens to the coast between cliffs up to 10 m high cut in stratified gravel (>Fig. 8.16.1.2). The adjacent beaches are of sand and shingle derived from the piedmont deposits, and there are boulders of basalt that have

Sardinia

fallen to the shore. Where the limestone outcrops in the lower cliff, there are notch and visor profiles, etched by solution as well as abrasion. North of Cala Gonone, the limestone cliffs return to the coast and continue to the Caletta di Cartoe, the mouth of a valley that descends from Dorgali. Beyond this, the Pleistocene basalt reaches the coast, forming a headland at Calleta di Osalia, and extending north past the long beach and dune-capped barrier spit at Orosei, where the Rio Codreto flows into the sea, and on from Punta Nera toward Cala Liberotto. Here, granite comes to the coast, and spurs of much-jointed brown granite run out between quartzose sandy beaches. To the north is a low plateau of weathered granite, incised by several valleys and backed by the barren Baronia Range. There are tors on rugged ridges that descend toward the coast, and at Capo Comino the shore below the derelict lighthouse has closely jointed granite disintegrating into rounded blocks. At Cape Comino, the coast swings westward and is low-lying, with ridges of granite running out into the sea between quartzose sandy beaches. These are backed by white dunes, partly held by scrub. Seagrasses (Posidonia) have been washed up on the beaches, and locally rounded into balls by wave swash and backwash. The sandy coast is scalloped in outline, with rocky protrusions that pass from granite northward to Palaeozoic metasediments. St. Lucia has dunes behind an irregular sandy beach strewn with brown seagrass hay on either side of the Rio Siniscola. Outcrops of grey schist and mudstone occur at intervals along the coast. Near La Caletta, the deltaic plain has rushy salt marsh and pines on low dunes behind the beach. At San Giovanni is a harbour on a coastal salient. To the north is a succession of sandy bay beaches between low headlands at Punta la Batteria, Punta dell’Asino, and Punta di Ottiola up to Teodoro. A sandy barrier encloses the Stagno di San Teodoro lagoon.

3.  The North Coast The southern boundary of the granite batholith of NE Sardinia comes to the coast north of Stagno di Teodoro, and extends past Punta Sabbatino and Capo Coda Cavallo, and on to the island of Molara to the north. Isola Tavolara and Capo Figari are of pale Jurassic dolomite. Isola Tavolara rises to 564 m and is fringed by tall cliffs with steep-sided inlets such as Cala di Levante and natural arches. Granite and schist alternate with small beachfringed lowlands in the Golfo di Olbia. Costa Smeralda consists of schist and granite, and to the north are many gulfs and inlets on a coast bordered by

8.16.1

the Arcipelago de la Maddalena, a group of high granitic islands with rocky shores and some beaches. Porto Cervo is a seaside resort with a large marina and to the west of Capo Ferro, a promontory of granite gneiss, are several inlets, with Baia Sardinia backed by a curving sandy beach. The funnel-shaped Golfo di Arzachena is the longest and widest of the gulfs opening northward, bordered by slopes that descend to sandy beaches. There are low promontories and nearshore sea tors of granite, and submerged calcareous ledges off beaches. Near Palau are the much weathered Bear Rock on Capo d’Orso, and a tombolo at Porto Polio. West of the Arcipelago de la Maddalena, the granite coast faces north across the Strait of Bonifacio toward Corsica. Punta Falcone is the northernmost point of mainland Sardinia. At Santa Teresa Gallura, a high terrace overlooks the granitic coast with Longone a tower on a promontory. To the west is Capo Testa, a prominent granite headland showing the response to variations in intensity of jointing and shearing, and weathering features including tafoni (>Fig.  8.16.1.3). There are small sandy coves. Pleistocene dune calcarenite occurs on the cliffs behind several of these (>Fig. 8.16.1.4). South from Capo Testa, the rocky granitic coast is backed by slopes mantled by rubbly drift, and there are sandy beaches along low-lying sectors, as at Spiaggia de Rena Maiore. The granitic shore is much dissected along joint planes and sheared zones, as on the Costa Paradiso (>Fig.  8.16.1.5) where the pink granite has rocky coves but no beaches, only artificial rocky hards and a launching ramp. Punta li Caneddi is a sharp promontory bordering the Baia Trinita, and to the south-west is Isola Rossa, a pink granite rocky island just offshore. The town of the same name has a large harbour protected by boulders, and small sandy beaches. The granite comes to an end at Badesi, SW of Isola Rossa, and there is an alluvial lowland, the Coghinas delta, fringed by a long sandy beach and dunes, interrupted at the river mouth. To the west of Valledoria Tertiary, sedimentary rocks and volcanic formations (Oligo-Miocene basalt lava and ash) outcrop along the coast to Punta Tramontana. Castelsardo is a town perched on a trachyte headland, its steep western coast descending to weathered red-brown rock outcropping along the shore. To the west, Lu Bagnu is a seaside resort with a beach backed by dunes, at the western end of which cliffs rise to 10 m, cut in Tertiary sandy limestone. In Peruledda Bay, the cliff is mantled by Pleistocene dune calcarenite above limestone ledges on the shore and a beach of dark cobbles of volcanic origin. To the west, the Rio Romasinu flows out over a flat limestone ledge underlain by clay, and there are tilted ledges of broken limestone.

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8.16.1

Sardinia

⊡⊡ Fig. 8.16.1.3 Granite dissected along joints at Capo Testa, showing tafoni. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.1.4 Cove near Capo Testa, showing dune calcarenite on the cliff. (Courtesy Geostudies.)

The cliffs decline past Punta Tramontana, and the coast is beach-fringed west of Maritza. The pale sandy beach is backed by low dunes covered with pine woods, and at Marina de Sorso the beach is backed by a low cliff where the dunes have been cut back to form cliffs up to

2 m high in earthy gravel. There are nearshore bars of sand and gravel. Further west, at Platamona Lido, the dune ­barrier is backed by a shallow lagoon, often dry. There are remnants of a Pleistocene inner barrier enclosing older lagoons with freshwater Tyrrhenian Limestone (Guilcher

Sardinia

1983; Ulzega and Ozer 1980). To the west, the dunes end as a fault brings up Upper Tertiary marine calcareous sediment, exposed in low cliffs along to Porto Torres. Cliffs up to 10 m high are cut in northward dipping Miocene limestone, with sectors of shore bench above sea level. There are outcrops of dry sandy nodular limestone, with a capping of Quaternary sand and gravel. Porto Torres has a large harbour (with ferries to Genoa) and a marina. To the west is a wide bay fringed by sandy beaches, which in some sectors are barriers fronting lagoons (some reclaimed). As the coastline curves northwestward, a sandy barrier encloses Holocene lagoons, Stagno di Pilo, and Stagno di Casaraccio, and again there are segments of an inner (Pleistocene) barrier impounding older lagoons (Guilcher 1983; Ulzega and Ozer 1980). Stagno di Casaraccio is a barrier lagoon with a sluice at an outlet opening to a sandy shore with rush swamp beside a former salt works, Tonnara Saline. It has low-lying shores with narrow beaches of gravel and sand, and some salt marsh. To the north, cliffs cut in Silurian shales and limestones extend past Stintino and out to Punta Negra. A north-facing bay is lined by Spiaggia della Pelosa, a sandy beach with dunes and a cuspate spit behind a shallow sandy area. The beach faces across a strait (Rada dei Fornelli) to Isola Piana, a long low scrubby island, with the larger and higher Isola Asinara massif beyond. This has generally rocky shores on Palaeozoic formations, with a granite waist south of Tumbarino. The Spiaggia della Pelosa runs NW to a point (Capo del Falcone) with limestone cliffs, and Torra Pelosa stands on a small island.

3.  The West Coast From Capo del Falcone, the west coast of Sardinia runs southward past Capo dell’ Argentiera, where Lower Palaeozoic phyllites outcrop, and there are Permo-Triassic rocks on either side of Porto Ferro. On the northern side is a steep coast on a limestone massif, and at the head of the bay a curving sandy beach backed by dunes and a lagoon (Lago Baratz). On the southern side, a thin dune calcarenite caps cliffs of grey limestone with structural spurs dissected on the shore. Torre Bantine Sala stands on the headland. Cliffs continue southward to Porticciola, where a small promontory has horizontal strata faulted against ­southward-dipping sandstone and red clay. This borders Cala Viola, a bay of clear water over rocks and seaweed with patches of sand, backed by a sandy bay-head beach,

8.16.1

but no dunes (>Fig. 8.16.1.6). On the southern side is a cliffed headland of conglomeratic sandstone and to the south the sandstones and clays dip seaward, until they are cut off by a fault against the high Jurassic and Cretaceous limestone massif with its precipitous cliffs extending towards Capo Caccia. The cliffs that border the Tramariglio Peninsula include escarpment cliffs over 100 m high fringing Punta Cristallo. They are plunging cliffs, with deep water near the shore (>Fig. 8.16.1.7). To the south, they are backed by gentler eastward slopes. The cliffs are precipitous above Cala dell’Inferno, where there is a high limestone island, Isola Foradada, and some dangerous rocky islets. In places, rock masses have fallen from the cliff leaving behind large concavities. The Jurassic limestone surface has thin brown clay soil, locally lowered into soil pipes. On the eastern side, there is the smaller Cala Calcina. Toward Capo Caccia is an isthmus with deep calas on both sides, and steps leading down to Grotta di Nettuno, famous for its dripstone formations. Below the steep massive limestone ⊡⊡ Fig. 8.16.1.5 The influence of jointing in the granite at Costa Paradiso. (Courtesy Geostudies.)

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⊡⊡ Fig. 8.16.1.6 The Torre Bantine Sala promontory at Cala Viola. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.1.7 High cliffs at Cala dell’ Inferno. (Courtesy Geostudies.)

are heaps of fallen rocks. Cape Caccia is crowned by a lighthouse in a military reservation. East of the Tramariglio Peninsula slopes descend to the bay of Porto Conte, which has low rocky shores with a prominent limestone ledge and only minor beaches. At its head are beaches and the remains of a Roman Villa on a coastal terrace about 2 m above sea level, fronted by a rocky ledge. The eastern coast of Porto Conte rises past Torre Nuova to an overhanging high cliff at Punta di Giglio. Then, the limestone cliffs face south along to Capo Galera. In the Bay of Alghero, there are sandy beaches be­­ tween low rocky headlands. Bombarda Beach is backed by pines on sandy ridges, and at Fertilia a coastal barrier with a pinewood is backed by a lagoon that has an outlet to the sea. The coast curves southward to Alghero, where the Old Town stands on a walled promontory. In places, the shore rocks have been excavated to leave an artificial platform. At Calabona, a southern coastal suburb, the shore has northward-dipping strata planed off and overlain by dune calcarenite, and there are small sandy beaches. South of Alghero, the coast steepens, with stratified limestones above shales that form irregular lower slopes. The coast road curves in round incised valleys. To the south, the land becomes higher, with many volcanic outcrops, and there are long coastal slopes incised by valleys. Locally, there are ubtidal algal benches. At Cabo Marargiu, the coast swings eastward, then curves round to Bosa, at

Sardinia

the mouth of the broad Tema valley. There is a sandy beach and a marina. The steep coast then resumes southward to Capo Nieddu, incised by the valleys of rivers such as the Rio Mannu. At Santa Caterina di Pittinuri is a cove with a cobble and sand beach at the mouth of an often dry river. The cliffs on either side are in white rubbly limestone capped by a slope of sandy clay. Spiagge dell’arco is a natural arch in the white bedded limestone south of this beach, and there is a group of white islets in the sea. To the south is the long sandy beach of Arenas, facing NW and backed by dunes carrying a pine forest. This ends at the low cliffed Tertiary limestone promontory of Capo Mannu, which is backed by lagoons and saline flats. On its southern side, the bay of Cala Marine is fringed by a sandy beach and barrier enclosing a lagoon at Putzu Idu. The beach is generally strewn with seagrass (Posidonia). To the south is the Sinis Peninsula. On its west coast, low cliffs of Tertiary limestone continue, with intermit­­ tent beaches, to Capo sa Sturraggia. At Mari Ermi, there is a long beach of quartzose sand fronting a swale lagoon, which is often dry. The quartzose sand has come from the granitic Isola di Mal di Ventre, several kilometres offshore. Further south is the beach of San Givanni di Sinis, which faces SW, and a narrow isthmus links Capo San Marco, an andesite headland that protrudes southward beside the Golfo di Oristano (>Fig. 8.16.1.8). East of Capo San Marco are the ruins of Tharros, an ancient town on the south-facing coastal slope. It was occupied by Phoenicians in the eighth century bc, and by

⊡⊡ Fig. 8.16.1.8 Capo San Marco. (Courtesy Geostudies.)

8.16.1

Romans from 238 bc, but was abandoned around 1070 and excavated in the nineteenth century. It may once have extended beneath the sea, for there are rectilinear patterns on the sea floor, though these may have been related to harbour foundations. A sandy beach on the north shore of the Golfo di Oristano borders an outer barrier, and there is an inner barrier enclosing lagoons, the Stagno di Mistras and the Stagno di Cabras, where Cabras is a fishing port with a long history on the eastern shore. Oristano stands about 4 km inland, south of the Rio Tirso. South of its harbour is a long sandy barrier beach on the east coast of the Golfi di Oristano, enclosing shallow lagoons bordered by salt marsh and reed and rush swamp. To landward is the plain of Arborea, a large irrigated farming area. The beach is of fine grey calcareous sand with Posidonia piled along the water’s edge and a litter of hundreds of rolled seagrass balls. The wide low sandy barrier has pine plantations. The Golfi di Oristano is framed by volcanic (andesite) headlands, Capo San Marco, and Capo della Frasca. The beach ends southward a longshore spit at Punta Corru Manu diverging from an inner spit that ends at a castello near Marceddi, beside the mouth of the Rio Mogora. The river mouth is enclosed by a dam with gaps through which the tide enters and leaves. The southern shores of the estuary had been cliffed before delta growth cut them off from the sea. The Rio Mogora drains a broad valley that drains part of fault-bounded lowland corridor (the Campidano rift valley) that runs SSE

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Sardinia

from here across to Cagliari, floored with Quaternary sediment. West of the Rio Mogora mouth, a hilly promontory runs out northward to Capo della Frasca, consisting of Upper Tertiary andesite. On its western coast, rugged cliffs extend south to Punta di s’Aschivoni, which has a flat lava flow and ends in an outlying large stack (s’Aschivoni). The volcanic rocks give place to Palaeozoic (Devonian) schists behind the Golfo di Porto Pistis. The coastal slope in slaty drift descends to a shore with dissected and potholed dune calcarenite and beach rock and beach conglomerate dipping seaward (>Fig. 8.16.1.9). To the south is a large dune extending inland, a site where dune calcarenite may still be forming. The Palaeozoic rocks extend along the Costa Verde, south to Capo Pecora, a granite headland. At Buggeru, the Rio Mannu outlet is diverted by a sand spit. South from Portixeddu Cambrian carbonaceous rocks, including an­­ cient limestones, come to the coast, forming bold cliffs at Masua. The Sugarloaf is a steep-sided outlying islet. Then a lowland backs the Golfo di Gonnesa. More Upper Tertiary volcanic rocks outcrop out to Capo Altano o Giordana and round to Portocuso. They include rhyolites and rhyodacites on Isola di San Pietro and Isola di Sant Antioco. There are stacks at Punta delle Colonne on the south coast of San Pietro, and the Isola di Sant Antioco is linked to mainland Sardinia by a tombolo,

with Stagno di Santa Caterina on the southern side. There are several spits. South of the Ponti is the Golfo di Palmas and the Palmas River delta. On the hilly southwest coast, there are more small ­barrier lagoons, notably Stagno de is Brebeis, near Punta Menga, which is enclosed by a curved Holocene outer ­barrier and split by the remnants of an older (Pleistocene) barrier, with lagoons on either side (Guilcher 1983). There  are headlands of Palaeozoic limestone bordering Cala su Truccu, and south of Punta Menga the beach in Porto Pino faces SW and is backed by dunes 10 m high on  a barrier enclosing the lagoon Stagnodeis Brebeis, which contains remnants of an inner barrier. There is granite bordering Punta di Cala Piombo, and Cambrian carbonaceous rock on Cape Teulada and round to Punta di Pirasu. Another granite sector occurs near Capo Malfatano. Capo Spartivento is also of granite, and as the  coast turns toward the Golfo di Cagliari there are more  beach-fringed lowlands, as at Pula, which has a ­tombolo (Capo di Pula). Several volcanic hills dot the coastal plain, and south of Sarroch an andesite volcano forms a hill with Torre del Diavolo on its eastern side. Sarroch overlooks the bay and a 9 km long sandy barrier borders the Stagno di Cagliari, which is also divided by remnants of a Pleistocene inner barrier (Guilcher 1983). This circuit of the coast of Sardinia is completed at Cagliari.

⊡⊡ Fig. 8.16.1.9 Dissected dune calcarenite on the shore of the Golfo di Porto Pistis, looking north towards the Punta di s’Aschivoni. (Courtesy Geostudies.)

Sardinia

References Cocozza T, Jacobacci A (1975) Geological outline of Sardinia. In: Squyres C (ed) Geology of Italy. Earth Sci Soc, Libyan Arab Repub, pp 49–81 Guilcher A (1983) Twin barrier-lagoon systems in Sardinia. In: Bird ECF, Fabbri P (eds) Coastal Problems in the Mediterranean Sea. Commission on the Coastal Environment, Bologna, pp 77–81

8.16.1

Ulzega WR, Ozer A (1980) Excursion - Table Ronde, Tyrrhénien de Sardaigne. INQUA Shorelines Commission, Cagliari Arba P, Arisci A, De Waele J, Di Gregorio F, Ferrara C, Follesa R, Piras G, Pranzini E (2002) Environmental impacts of artificial nourishment of the beaches of Cala Gonone (Central-East Sardinia). Proceedings of Littoral 2002, Eurocoast Portugal 465–468

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8.16.2 Sicily

Paolo Ciavola

1.  The Coastline of Sicily At the NE corner of Sicily Capo Peloro is an urbanised spit, with a beacon tower. The north coast is mountainous, a westward continuation of the Calabrian Appennines with sandy and gravelly beaches occupying wide embayments. Much sand and gravel is delivered to the coast by rivers in incised valleys. There is evidence of Holocene uplift (Rust and Kershaw 2000). West of Cape Orlando long sand and gravel beaches extend to the Cefalù headland. Termini Imerese has a large industrial area with an oil and chemical terminal and an alternation of small pocket beaches surrounded by cliffs along to the city of Palermo. This part of the coast is heavily urbanised. The beaches east to the Golfo del Castellamare are long and of sandy nature, with the exception of Guidaloca Beach, formed by white cobbles. The large gulf of

Castellamare is delimited on the western side by the steep slopes of the Natural Reserve of the Zingaro, an area with high biodiversity and with many endemic species of plants. Capo San Vito at the NW corner of Sicily is followed south–westward by the beach of San Vito lo Capo, which has an important calcareous fraction due to in situ biogenic production. The calcareous nature of the landscape has fa­­voured the formation of karstic caves and underground drain­­age. To the south is the Monte Cofano massif (>Fig. 8.16.2.1), with an height of more than 600 m. formed by marble that has long been quarried, then a coastal plain to Trapani. Off Trapani there is a small archipelago (Isole Egadi) formed by Marettimo, Favignana and Levanzo. Marettimo is located further away offshore and has a central mountain that raises more than 600 m above sea level. Dolomite cliffs (>Fig. 8.16.2.2) are present all round the island and beaches are absent.

⊡⊡ Fig. 8.16.2.1 View of Monte Cofano.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.16.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.16.2.2 The dolomitic limestones of Marettimo Island.

⊡⊡ Fig. 8.16.2.3 Windmills and salt pans near Mozia.

The coastline just south of Trapani enclosed a large wetland (Stagnone di Marsala) with an island at its western edge. The whole landscape is dominated by windmills and saltpans testifying a long tradition of salt production (> Fig. 8.16.2.3).

Along the south coast, facing the Sicilian Channel, beaches are more extensive, fronting a narrow coastal plain punctuated by hilly promontories. There are emerged Pleistocene marine terraces in the hinterland, and the offshore profile declines steeply seaward beneath the Canale

Sicily        

di Sicilia. Wave energy here is high, and well developed dune systems are present along the coast of Ragusa, eastward to Santa Maria del Focallo. Near Realmonte limestones and marls form coastal cliffs with a step-like outline (Scala dei Turchi). A succession of promontories and bays with meagre beaches extends further south to Capo Passero, where the Isola delle Correnti marks the boundary between the Ionian Sea and the Sicily Channel. The point is the furthermost tip of the island, with a latitude equivalent to that of Tunis. Around the cape beaches are adjacent to wellpreserved dune systems. Notable is the Vendicari coastal lake system (>Fig. 8.16.2.4), a nature reserve of Ramsar interest, managed by the Sicilian regional authorities since 1989. The lakes are part of a group of wetlands that extends for most of the coast north to Syracuse. Many of these have silted up or have been reclaimed (Geremia et al. 2004), but Vendicari is still virtually untouched. The wetlands are separated from the open sea by a dune ridge 2.5 km long, interrupted by an inlet that in the past was active in the winter season but that at present is occupied by embryo-dunes and a wind-swept storm beach (Ciavola et al. 2005). A small rocky island is present at centre of the barrier island system, forming a tombolo surrounded by Posidonia meadows. A succession of promontories and bays with large pocket beaches extends north from the Torre di Vendicari.

⊡⊡ Fig. 8.16.2.4 The Vendicari coastal wetland, used for centuries as salt pans. The low dune barrier is in some sectors ­occasionally overwashed. On the foreshore is debris of Posidonia that forms extensive meadows on the nearby seabed.

8.16. 2

A high cliff in limestone can be seen at Capo Murro di Porco. The Ortygia island off Syracuse was artificially linked to the mainland in the fourth century bc and later became one of the main parts of the Greek town. There are cliffs cut into marls at Punta Castelluccio. To the north is the Golfo di Augusta, which contains a number of ports used since Greek times and largely developed in the sixteenth and seventeenth centuries during the Spanish and French domination. At the southern edge of the bay there is the Peninsula Magnisi, which unfortunately is adjacent to the chemical terminal of Priolo, with huge pollution problems. The coastline north of Augusta continues in an alternation of cliffs and small beaches, an area rich in saltpans which once were used for salt production. The wide beach (Piana di Catania) bordering the lowland behind the Golfo di Catania is subject to occasionally strong easterly wave action. Beaches are formed of fine sand and have dissipative nature with systems of multiple nearshore bars. Wave energy diminishes northward as the Ionian Sea narrows to the Strait of Messina. The sediment input into the area is provided by the Simeto River, the largest river in Sicily for what concerns water discharge and size of the catchment, which flows around the southern side of Etna. The present deltaic area has been transformed by man into agricultural land through reclamation to fight malaria around The Second World War.

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⊡⊡ Fig. 8.16.2.5 Basaltic lavas and sea stack at Isola Lachea in front of Acitrezza. The site is a marine reserve.

⊡⊡ Fig. 8.16.2.6 The harbour at Giardini Naxos on the east coast of Sicily. Sand has ­accumulated in the lee of the breakwater. (Courtesy Geostudies.)

Sicily        

⊡⊡ Fig. 8.16.2.7 Beach rock on the shore at Giardini Naxos. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.2.8 The Taormina promontory. (Courtesy Geostudies.)

8.16. 2

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⊡⊡ Fig. 8.16.2.9 The limestone cliff at Capo Taormina. (Courtesy Geostudies.)

⊡⊡ Fig. 8.16.2.10 Tombolo of Isola Bella north of Capo Taormina. (Courtesy Geostudies)

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8.16. 2

⊡⊡ Fig. 8.16.2.11 The steep coast north of Taormina. (Courtesy Geostudies.)

From Catania north to Pozzilo lava flows reach the coast, forming high cliffs. Around Acireale the cliffs reach the highest point in the Timpa area, which has developed a unique ecosystem related to the freshwater outflows at the base of the cliffs. At Acitrezza there is a small group of sea stacks (faraglioni) which is believed by Greek mythology (Odyssey) to correspond to the stones thrown by blind Polifemo to Ulysses (>Fig. 8.16.2.5). North to Risposto the beaches are narrow and backed by a small coastal cliff cut into sedimentary deposits (Chiancone Formation) formed by weathering of previous volcanic cones of Mount Etna. Inland there is indeed the intermittently active volcano, Mount Etna, which has lava flows down to the coast. There was a major eruption of Mount Etna in 1699, and many subsequent eruptions have augmented the volcanic cone since. There is a continuous long sandy and gravel beach north of Risposto, and near Capo Schisò the outflow of the Alcanta River forms a small delta after cutting a deep canyon (Gole dell’Alcantara) into columnar basalts. At Giardini Naxos the beach has drifted southward in the lee of a harbour breakwater, and rocky groynes at the northern end have failed to retain it (>Fig.  8.16.2.6). Beach rock is exposed on the shore where the beach has been depleted (>Fig. 8.16.2.7). North of the bay of Giardini is Taormina, where a ­calcareous mountainous range (>Fig.  8.16.2.8) descends eastward to the coast. The Taormina range ends in a bold headland with steep limestone cliffs at Capo Taormina

(>Fig. 8.16.2.9). A tombolo (>Fig. 8.16.2.10) attaches to a small rocky island here (Isola Bella). To the north sand and gravel beaches, fed by gravelly watercourses, extend in front of a narrow coastal plain, interrupted by steep rocky headlands (>Fig.  8.16.2.11). There are a number of beach compartments all the way up to Messina. On the western side of the Strait of Messina there is a system of small coastal lakes and spits at the base of Cape Tindari (>Fig. 8.16.2.12), which in recent time has been affected by shore protection works (Randazzo et al. 2005). The beaches continue to a recurved spit that protects the harbour of Messina. It is the outcome of northward longshore drifting on the east coast bordering the deep Strait of Messina. The circuit of Sicily is completed at Capo Peloro.

2.  The Eolie Islands Volcanic landforms occur on several islands north of Sicily (The Eolie Islands), notably in Lipari, where marine processes have dissected volcanic (andesite) structures, and Stromboli and Vulcano are intermittently active. Here,  too, there is evidence of tectonic displacement (Calanchi et al. 2002), which has also been traced across many other points around the island using quaternary indicators (Antonioli et al. 2006). An island off Stromboli rose out of the sea then sank again in 1955. The Lipari Islands also include Lipari, Salina, Panarea, Alcudi and

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⊡⊡ Fig. 8.16.2.12 The Tindari headland spit.

⊡⊡ Fig. 8.16.2.13 Steep coast on volcanic rocks, Filicudi, Lipari Islands.

Sicily        

8.16. 2

⊡⊡ Fig. 8.16.2.14 Rugged coast along the western shore of the island of Pantelleria.

Filicudi. They have steep scrubby ravined slopes and dark sand and gravel beaches derived from the volcanic rocks (>Fig. 8.16.2.13).

3.  Pantelleria and Lampedusa Finally, Sicily has two small islands located in the middle of the Canale di Sicilia, closer to Africa than to mainland Italy. Pantelleria (>Fig. 8.16.2.14) is located over 100 km south of Sicily and has a volcanic nature, with the central vulcano rising over 800 m above sea level, which is inactive since 1891. The island has no beaches but many coves and coastal caves. A small salt lake is present in the inner part. Lampedusa is a bare, low-relief island, 11 km long, with the largest beach of the Rabbit Island (Isola dei Conigli) where turtles lay their eggs, one of the few remaining sites in the Mediterranean. Linosa is the smallest of the Pelagic Island, a volcano which has been inactive since 2,000 years ago, with rocky coves and black sand beaches.

References Antonioli E, Kershaw S, Renda P, Rust D, Belluomini G, Cerasoli M, Radtke U, Silenzi S (2006) Elevation of the last interglacial highstand in Sicily (Italy): a benchmark of coastal tectonics. Quatern Int 145–146:3–18 Calanchi N, Tranne CA, Radtke U (2002) Late Quaternary relative sea level changes and vertical movements at Lipari, Aeolian Islands. J Quatern Sci 17:459–467 Ciavola P, Armaroli C, Balouin Y, Geremia F (2005) The back-barrier wetland system of Vendicari (SR): prediction of dune overwash­­ ing  using shoreline variability indicators. Proceedings of  Lagoons and Coastal Wetlands, ICAM dossier n. 3, UNESCO, Venice, pp 307–314 Geremia F, Lanza S, Randazzo G (2004) Le zone umide costiere in Sicilia: individuazione, classificazione e valutazione. Atti dei convegni dei Lincei 205:215–223 Randazzo G, Geremia F, Lanza S (2005) Negative response to remedial measures of shore protection in Sicily: the case of the Tindari headland spit (Northern Sicily). Eurocoast Portugal 199–200 Rust D, Kershaw S (2000) Holocene tectonic uplift patterns in northeastern Sicily: evidence from marine notches in coastal outcrops. Mar Geol 167:105–126

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8.17 Malta

George Said · John Schembri

1.  Introduction The Maltese Islands (notably Malta, Gozo, and Comino) are a group of central Mediterranean islands located about 96 km from Sicily and 290 km from North Africa. They are dominated by limestone formations, and much of their coastline consists of steep or vertical limestone cliffs, indented by bays, inlets, and cliffy coves (>Figs. 8.17.1 and > 8.17.2). The climate is Mediterranean, with hot dry summers and

cool rainy winters: Valletta has mean monthly ­temperatures of 12.8°C in January and 25.6°C in July, with an average annual rainfall of 578 mm. Westerly winds are prevalent, but waves can be generated from all directions. The tide range is small, attaining a maximum of about 40 cm. The islands rest on the submerged Malta–Hyblean platform, a wide undersea shelf bridge that connects the Ragusa platform of southern Sicily with the Tripolitana platform of northern Libya. They occupy the zone of

⊡⊡ Fig. 8.17.1 A map of the Island of Malta displaying the main coastal geomorphological features.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.17, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Malta

c­ onvergence of the African and Eurasian plates, and tectonic movements have tilted and faulted the rock formations, which are of mid-Tertiary age (Hyde 1955). The following formations are present in descending order (Pedley et al. 1976): •• •• •• •• ••

Upper Coralline Limestone (up to 162 m) Greensand (0–11 m) Blue Clay (0–65 m) Globigerina Limestone (23–207 m) Lower Coralline Limestone (at least 140 m)

The Upper Coralline Limestone is a resistant stratified limestone of Miocene age, which generally forms vertical cliffs rising to a broad, gently undulating plateau or to residual mesas. The underlying Greensand, Blue Clay, and Globigerina Limestone are softer formations forming more gently sloping coasts, and excavated as bays and inlets where they outcrop at and below sea level. The Greensand, a glauconitic bioclastic formation, is several

metres thick on Gozo, but less than a metre on Malta: it weathers to orange brown sand that is the source of several beaches. Concave slopes are found on Blue Clay outcrops, and shore platforms have developed on resistant layers in the well-jointed and stratified Globigerina Limestone where waves have swept away overlying soft or weathered rock (Guilcher and Paskoff 1975). The Lower Coralline Limestone, of Oligocene age, is another resistant formation, which forms bold cliffs in western Gozo. The coast of the Maltese islands has been shaped by past and present tectonic activity, the lithology of outcropping rock formations and the effects of changing sea levels. Tectonic activity in the central Mediterranean created the land mass of the Maltese archipelago as an uplifted block, probably bordered by scarps on fault lines or monoclines, which have been cut back as cliffs. Malta has been tilted about 4° in a northerly direction, producing cliffs up to 200 m high on the south and southwestern coasts and slopes descending to low cliffs and rocky shores on the

⊡⊡ Fig. 8.17.2 A map of the Island of Gozo displaying the main coastal geomorphological features.

Malta 

northern and eastern coasts. Gozo has a gentle easterly dip, so that the Lower Coralline Limestone, which forms high cliffs on the west coast, declines below sea level, but it reappears on the east coast at Qala Point. On both islands, ridges and valleys have been shaped by denudation of the component rock formations by rivers and solution processes in patterns influenced by faults, which generally run E-W. The Victoria Lines escarpment has formed along a fault line that runs from west to east across northern Malta. It is one of a series of parallel faults that between uplifted horsts, as at Mellieha, and subsided grabens, valleys that open to coastal embayments and inlets, have been cut out between fault lines, as at St. Pauls Bay and Mellieha Bay. Movements have occurred on some of these faults during recurrent earthquakes, as in March 1972. The cliffs that border the Maltese Islands are not directly related to faulting, although some sectors, as in Fomm-ir-Rir Bay, follow fault lines, and the Dingli Cliffs pass eastward into a fault line scarp along the major WNWESE Maghlaq Fault, which has downthrown Upper Coralline Limestone southward by at least 230 m (Pedley et al. 1976). Vertical plunging cliffs, which sometimes have minor notches at present sea level and no shore platforms, descend directly to the sea floor and were probably formed when sea level was lower, relative to the land. Contrasts in lithology influence cliff profiles, the harder Upper and Lower Coralline Limestone forming vertical cliffs while the Greensand, Blue Clay, and Globigerina Limestone form debris-strewn coastal slopes that show evidence of  slumping, particularly where the Upper Coralline Limestone has collapsed over Blue Clay lubricated by percolating groundwater. Coastal slopes of this kind are known locally as Rdum cliffs. Cliff outlines have been strongly influenced by lithology and structure. Cliff recession occurs when crevices form roughly parallel to the cliff line, generally 10–20 m inland, and the rock mass to seaward subsides or slumps, disintegrating into boulder aprons. Gorges (geos) have been cut out along faults and joint planes that run perpendicular to the coastline, particularly in the Lower Coralline Limestone, and caves, natural arches, and stacks have been formed by marine erosion along joints and bedding planes in these cliffs. Similar features are seen on the bedded sedimentary Globigerina Limestone, where weathering and wave erosion have produced stacks, blowholes, and natural arches, as on the west coast of Gozo. Karst weathering is evident on rocky shores, cliff faces, and cliff crests where the unvegetated zone often extends well inland. Sea level changes accompanying climatic variations through Quaternary times have also influenced geomorphological evolution: about 18,000 years ago, in Late

   8.17

Pleistocene times, for example the sea stood about 120 m below its present level, and the Maltese archipelago was one large island. There are submerged shore platforms in Marsaxlokk Bay in Malta, along the Xwejni coast in Gozo, and 50–60 m below sea level beneath plunging cliffs. During low sea level phases of the Pleistocene, some valleys were excavated in what is now the sea floor, and the Late Quaternary (Flandrian) marine transgression, which brought the sea up to its present level about 6,000 years ago, invaded these to produce rias, as in Grand Harbour, Valletta, where rivers etched out a valley system that was later submerged by the sea to form inlets and harbours (>Fig. 8.17.3). Marine submergence of narrow incised valleys has produced calanques, as at Mgarr ix-Xini, where the vertical walls descend to a sandy submerged valley floor and Wied il-Ghasri on the north coast of Gozo. There has evidently been very little change on plunging cliffs during the 6,000 years since the sea arrived at its present level. Semicircular bays have formed on cliffed coasts where dolines have been invaded by the sea. Dolines are depressions in a limestone surface, generally circular or oval in plan, with a range of forms: dish or bowl shaped, conical, and cylindrical. They vary in size, and are often flatfloored, lined by impermeable clays, the insoluble residues of limestone consumed by solution processes as the land surface was lowered, as at Qawra in Gozo. The rounded bays at Xlendi and Dwejra on the west coast of Gozo originated as underground caverns with roofs that collapsed to form dolines that some were submerged by the sea. Shore platforms on the Maltese limestone are of three types. There are seaward sloping intertidal platforms produced mainly by abrasion, structural platforms with ledges related to resistant rock outcrops that are horizontal or gently dipping and solution platforms, sub-horizontal at low tide line. They have developed since the sea reached its present level about 6,000 years ago. Seaward sloping intertidal platforms are produced by wave quarrying and abrasion and are best found in areas of high wave energy and easily quarried rock such as the Globigerina Limestone. They occur on sectors of the coast exposed to strong wave action, such as the stormy northwest and the north coasts of Malta and Gozo. Structural shore platforms have developed where a horizontal or gently dipping resistant rock formation outcrops at sea level, as on the south coast of Gozo. Shore platforms produced by solution of the limestone, accompanied by repeated wetting and drying, are seen on the north and east coast of Malta. The irregular surface produced by ­solution and wetting and drying is well illustrated on the shore at Dwerja in western Gozo. Coastal slopes on the  Globigerina Limestone show stepped profiles, with

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⊡⊡ Fig. 8.17.3 The Grand and Marsamxett Harbours, a ria coast on the east of Malta. (Courtesy of the Department of Information Malta.)

s­ tructural benches on outcrops of thin hard pebble beds, ending in visors above notches cut in softer intervening beds. These stepped profiles are prominent near Marsalforn in Gozo and occur intermittently between Ricasoli and Delimara on Malta. Beaches are rare on the coast of the Maltese islands, making up less than 2% of the coastline. They do not form along vertical plunging cliffs because these reflect wave action and prevent deposition of sand or shingle. They are found in some bays and inlets, and are generally derived from the Greensand where it outcrops in cliffs, across the shore and on the shallow near shore sea floor; there is ­usually an admixture of calcareous biogenic sand formed from sea floor organisms and swept in by wave action. Dunes have formed behind a few wide beaches, as at Ramla on the north coast of Gozo. Shingle is derived from the gravelly material produced by the disintegration of thinly bedded or closely jointed outcrops of Globigerina and Lower Coralline Limestone along the northern and eastern coast of the islands, where waves are eroding low cliffs and shallow submerged rock outcrops.

2.  The coasts of the Maltese Islands 2.1.  Malta The city of Valletta stands on a limestone peninsula that runs out to St. Elmo Point (>Fig. 8.17.3). To the southeast is the Grand Harbour, and to the northwest Marsamxett

Harbour, both with bays and inlets known locally as Creeks. Northwest from Dragut Point, broad limestone slopes descend to St. George’s Bay and low rocky karstic shores with segments of shore platform, as in Bahar-icCaghaq Bay and on Qrejten Point, and solution basins as at Sliema. There has been extensive development, notably at Sliema and St. Julians. The low-lying coastline is fairly regular until interrupted by the wide Qalet Marku Bay, beyond which is the Ghallis Point Promontory, then valley-mouth Salina Bay, cut in soft Globigerina Lime­ stone, with its numerous salt pans. A long peninsula runs out to Qawra Point, then St. Pauls Bay is deeply inset, with parallel coasts and curving bay-head Pwales Beach. Cart tracks, possibly of Neolithic age, incised into limestone pass beneath the sea and reappear on the farther shore (Hyde 1955). The northern coast of St. Pauls Bay has some cliffy sectors, as at Rdum Irxaw, and out to Blata L-Bajda, with St. Pauls Island offshore. On the northern coast of this peninsula, bordering Mellieha Bay, there are bays and steep slopes. Mellieha Beach is a curving sand and gravel beach at the shallow head of the bay, with red-brown sand derived from weathered Greensand on the bordering shores and sea floor. The red sand outcrops in the corner of Marfa Point, with a small beach. The north coast has a succession of small coves, ­­Ghartal-Madonna and Tal-Imgharqa, a natural arch at Il-Parsott, then higher cliffs to Il Marbat and the headland of Dahlet Ix-Xilep. Here, the coast turns northwest and has high cliffs in Upper Coralline Limestone with tumbled rocks on the shore. The north coast is embayed with small

Malta 

­ eadlands including Marfa Point, which shelters the Gozo h ferry terminal. Paradise Bay, to the west, is followed by a sector of scree-covered slope beneath a limestone cliff, then the headland at Ras Il-Qammieh. The limestone cliff continues as the coast swings southward, and passes into a scarp along the northern side of the Mellieha valley. The western end of this valley is truncated by vertical cliffs in Upper Coralline Limestone, which rise to the point at Ras in-Niexfa. The cliff crest is backed by very rugged karstic limestone with enclaves of terra rossa clay. On the cliff face are breakaways and toppling columns in the Upper Coralline Limestone capping, and caves, boulders, and a notch at the cliff base. The plateau declines toward Anchor Bay, where there are long and deep arcuate crevices that form parallel to the cliff crest, and subsidence has occurred on the seaward side of these, with disintegrated blocks sliding down the coastal slope. On the northern side of Anchor Bay are the odd wooden structures built as Sweethaven, a film set for Popeye the Sailor made here in 1979. There is a small sandy bay-head beach, possibly derived from a yellow horizon in the adjacent cliffs. To the south are cliffed headlands and steep-sided bays out to Ras Il-Wahx. The cliffed coastline then recedes into Golden Bay, where the broad beach is of pale yellowbrown sand, similar in color to the Globerigina Limestone in the cliffs, and Ghajn Tuffiena Bay, which has a redbrown sandy beach derived from the Greensand, backed by subrounded limestone pebbles then a landslide slope on Blue Clay beneath the Greensand and a capping of Upper Coralline Limestone. This ends southward in a steep narrow promontory where an arête of Blue Clay runs out to a small mesa of Upper Coralline Limestone. To the  south is Gnejna Bay (>Fig. 8.17.4), with slopes in Blue  Clay descending to cliff-base ledges of Globerigina Limestone between more red-brown beaches. Beyond this the cliffs run out to Ras il-Pellegrin. The limestone cliff is fronted by a steep slope descending to Fomm Ir-Rim Bay, the southern side of which is a north-facing cliff along the fault zone of the Victoria Lines escarpment. This is one of the few sectors of cliff in Malta that coincides with a fault. It runs out to Ras ir-Raheb, where the coast turns south and a cliff in Upper Coralline Limestone is fronted by a steep slope of talus over Blue Clay, Rdum tal-Vigarju. The limestone cliff recedes behind a broad amphitheatre at Mtahleb, as the coastline turns southeastward. The Dingli coast consists of a karstic plateau ending in a vertical receding cliff of Upper Coralline Limestone (>Fig. 8.17.4), below which is a boulder-strewn slope of Blue Clay and Globigerina Limestone, ending abruptly in

   8.17

a ­vertical cliff on Lower Coralline Limestone, which plunges into the sea (>Fig. 8.17.5). There is a shore platform beneath the small promontory at Il-Kullana, below Ta’ Zuta, and to the east Fommir-Rih is a cove marking the beginning of a downfaulted foreland at Hagras-Sewda, which is backed by a steep slope in Lower Coralline Limestone overlain by Globigerina Limestone at TalBajjad. The downfaulted foreland consists of tilted Upper Coralline Limestone, downthrown at least 230 m by the Maghlak Fault. The steep backing slope is a fault line scarp along the Maghlak Fault, and there are exposures of the smooth fault plane with slickensides (Pedley et al. 1976). The spectacular vertical cliffs in Lower Coralline Limestone to the west are not fault scarps, however, for they have been cut back by marine erosion from the westward extension fault line scarp. To the east, the valley of Wied iz-Zurrieq descends to a narrow calanque, and the next cove contains the Blue Grotto, Wied Babu is another valley-mouth inlet, and the cliffed coast continues to Benghisa Point. This marks the beginning of the large Marsaxlokk Bay, cut In Globigerina Limestone and bordered by cliffs of Upper Globigerina Limestone that diminish behind Pretty Bay and St. Georges Bay, on either side of the Birzebbuga headland. Beaches have been artificially nourished here, using sand dredged from the harbour in Marsaxlokk Bay. Cart tracks incised into the limestone, possibly of Neolithic age, pass beneath St. Georges Bay (Hyde 1955). Marsaxlokk stands at the head of the bay, and cliffs increase along the eastern coast to Delimara Point. The east coast of Malta has a series of rounded bays separated by promontories including the long, narrow Ras il-Fenek, St. Thomas Bay, and Marsakala Bay. Low cliffs continue past Zonqor Point to Ricolosi Point on the eastern side of the entrance to Grand Harbour at Valletta.

2.2.  Comino North of Malta, the Comino Channel separates Gozo, and within it is Comino Island, consisting of Upper Coralline Limestone. Marine erosion has bitten deeply into the west coast, excavating the Blue Lagoon between Comino and the smaller western island of Cominotto, but the north and south coasts slope gently into the sea, the lower slope dissected by karstic weathering from spray and occasional storm swash and rainfall, a declining platform passing to a submerged rocky surface or terminating in a low cliff. There are higher cliffs on the northeast coast. Several bays have been cut out along joints.

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⊡⊡ Fig. 8.17.4 Sloping cliffs in Blue Clay at Gnejna Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 8.17.5 The rubble-strewn slope of Globigerina Limestone at Dingli on the south coast of Malta ends in vertical cliffs cut in Lower Coralline Limestone. A structural bench is seen at the base of these cliffs. (Courtesy Geostudies.)

2.3.  Gozo Gozo is reached from Malta by way of Mgarr Harbour, protected by a breakwater in a bay on the south coast. Slopes of talus-mantled Blue Clay descend to low cliffs in Globigerina Limestone eastward to Qala Point, the easternmost point of Gozo, where there are boulder cliffs in the underlying Lower Coralline Limestone. There is a sea­ ward sloping shore platform with a pinnacled surface and

solution pools. Above the cliffs, a long rubbly slope rises through Globigerina Limestone and Blue Clay to a capping of Upper Coralline Limestone on a mesa at Nadur. Steep slopes and basal cliffs run round to Dahlet Qorrot Bay, beyond which are Mistra Rocks, a lobate headland bearing scree below an Upper Coralline Lime­ stone cliff with scree mantling a slope that descends to the shore on Globigerina Limestone, the underlying Lower Coralline Limestone having passed below sea level.

Malta 

San Blas Bay occupies a valley mouth, and Rdum il-Kbir is another lobate headland crowned by a limestone cliff. The deep Ramla valley descends north to Ramla Bay, which has a sandy beach backed by sparsely grassed dunes, and a gentle seaward sloping, smooth shore platform. Rdum tax-Xaghra is a promontory at the end of the next ridge. The Marsalforn valley then descends to Marsalforn Bay, and to the west are coastal ledges on resistant layers in the Globigerina Limestone and segments of sloping coast on outcrops of Blue Clay. Wied Il-Ghasri is an incised valley, the lower part of which is submerged by the sea as a calanque (>Fig. 8.17.6). To the west, headlands and bays extend to Xwejni Bay, and a steep slope runs out to Reqqa Point. Here, the Lower Coralline Limestone re-emerges, and soon forms high cliffs. These are trenched by a long, narrow calanque, Ghar-il Qamh, and diversified by capes at Forna Point and Pinu Point. Wied-il-Mielah is another calanque between valley-side limestone cliffs. Beyond Hekka Point, high plunging cliffs of Lower Coralline Limestone extend to San Dimitri Point, rising 90 m above sea level, where the coastline turns southward. Dolines (Miocene solution subsidence basins in the Lower Coralline Limestone that were infilled with Globigerina Limestone) have been excavated to form rounded bays, and at Dwerja a lagoon (the Inland Sea), fringed by a curving beach of well-rounded pebbles, occupies a breached doline, and is linked to the sea through a cave, Tieqa Zerqua or the Azure Window, in the rim to the north, extending out to Dwerja Point.

⊡⊡ Fig. 8.17.6 The calanque at Wied i Ghasri on the north coast of Gozo. (Courtesy Geostudies.)

   8.17

Dwerja Bay is another doline, invaded by the sea through breaches on either side of a high rim remnant, Fungus Rock. The rocky backshore on Lower Coralline Limestone has been made irregular by solution and wetting and drying processes, while the overlying yellow Globigerina Limestone forms smooth surfaces. Down on the shore is a segment of flat solution platform below a notch on the soft Globigerina. Limestone Cliffs in Lower Coralline Limestone continue south to Wardija Point, then east along the Gebel Ben Gorg on the south coast of Gozo. Xlendi Bay is a slot inlet at the mouth of the long incised Xlendi valley, with another cove on the southern side where the cliffs run out to Ras il-Bajda. To the east, the limestone plateau ends in a cliffy embayed coast south of Sannat, extending past Ta’Cenc. Two incised valleys, Wied Sabbar and Wied Hanzira, converge in the branched calanque of Mgarr ix-Xini. There is a sand and gravel beach at the calanque head (>Fig. 8.17.7) and the bordering limestone cliffs have a notch at sea level. The Lower Coralline Island dips beneath the Globigerina Limestone here, but persists in the outlying Fessej Rock. Ras il-Hobz is a small rounded headland, and the cliffs decline to steep coastal slopes eastward to Ix-Xatt L-Ahmar, where there are slopes in Blue Clay descending to shore ledges on Globigerina Limestone (>Fig. 8.17.8) and stepped profiles within the Globigerina Limestone shoreline. Below Fort Chambray, the steep slopes of Tafal extend round a point to the Mgarr Harbour.

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⊡⊡ Fig. 8.17.7 The sandy beach at the head of the Mgarr ix-Xini calanque. (Courtesy Geostudies.)

⊡⊡ Fig. 8.17.8 Structural and stepped shore platform on Ix-Xatt L-Ahmar, south coast of Gozo. (Courtesy Geostudies.)

2.4.  Artificial Coastline The man-made (artificial, anthropogenic) part of the coastline of the Maltese Islands approximates to about 21% of their total length. No part of this type of coastline has been constructed for erosion control but mainly developed as part of port, recreational, and urban infrastructure. Most of these utilities are present in locations where

the coast is the most accessible. As a result, practically all  the concrete platforms, jetties, breakwaters, promenades, and slipways are to be found either in the densely populated localities such as the Grand and Marsamxett Harbour areas or at Marsaxlokk Bay where a huge container terminal, an oil blending facility, and six fuel storage depots are located, together with an increasing urban population.

Malta 

Touristic development and the enhancing of coastal areas with shoreline modifications to accommodate beach facilities can be found in St. Paul’s Bay, Mellieha, and the Sliema to St. Julian’s areas. Changes to the coast in Gozo are minimal. These are seen at Mgarr Harbour, Xlendi, and Marsalforn. As a result of these modifications, there are areas where the beach size has decreased and sediment quality deteriorated. Future sea level changes are not expected to improve the problem if structures to combat coastal erosion are not developed.

   8.17

References Guilcher A, Paskoff R (1975) Remarques sur la géomorphologie littorale de l’archipel maltais. Bull Assoc Géog Français 427:225–231 Hyde HPT (1955) The geology of the Maltese Islands. Lux Press, Malta Pedley HM, House MR, Waugh B (1976) The geology of Malta and Gozo. Proc Geol Assoc 87:325–341 Pedley M, Huges M, Clarke PG (2002) Limestone Isles in a Crsytal Sea: The geology of the Maltese Islands. Publishers Enterprises Group, Malta

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8.18 Slovenia

Roger Charlier

groundwater, and also by sea water and spray. Solution processes and wave action produce notches, deep narrow indentations at the base of the cliff, which then overhangs. Solution also produces pipes, pans, pits, and lapies on the shore, lapies being surfaces covered by a network of furrows and crests. Water temperatures in the Adriatic Sea range from about 12°C in February to 25°C in August. Warm water flows in from the Mediterranean Sea, so that even during harsh winters the Adriatic Sea is never colder than 10°C. The mean salinity of the sea is 3.58%, decreasing near river mouths: it is lower in May and December, when river discharge is larger, and higher in February and September. Tide range is very small (generally less than 50 cm), but strong winds from the south and west over the Adriatic Sea can raise water levels and generate waves 2–3 m high on the coast.

Slovenia’s (Slovenija) coast lies on the Istrian Peninsula on the NE coast of the Adriatic Sea. It extends some 100 km from the Gulf of Trieste to the Gulf of Kvarner, also known as the Kvarner Narrows. Istria has a mostly limestone coast with some areas of Tertiary flysch (highly fissile sandy and calcareous shales with occasional beds of conglomerates and/ or sandstone). Beaches are rare. The flysch is cut by ­elongated inlets (canali), typical of the Dalmatian type of coast, which often provide excellent harbour sites. They were formed by the partial submergence of valleys in a coastal region of parallel anticlines and synclines, which also separated numerous long and narrow high islands (>Fig. 8.18.1). The cliffs and steep slopes bordering the canali show karstic weathering (Cvijič 1924). This coastal karst forms where the rock outcrops are dissolved by precipitation and ⊡⊡ Fig. 8.18.1 Part of the Dalmatian coastline.

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.18, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Slovenia

The shape of this coast is so distinctive that geomorphologists recognise a “Dalmatian coast” as a separate type, a drowned mountainous coast that consists of ranges and straits that follow the axes of parallel folding. Dalmatian coasts are steep coasts with a longitudinal structure, producing elongated islands and peninsulas separated by straits, channels, and long, narrow gulfs (De Martonne 1948). Shepard (1973) classified ­Dalmatian-type coasts as a kind of ria coasts, as was also suggested by Baulig (1930) and von Richthofen (1885). Occasionally these inlets are referred to as fjords (e.g., the fjord of Kotor), but this is incorrect, as they were not produced by glaciation. The Dinaric Alps end in steep slopes overlook­­ing narrow coastal plains, or descending straight to the sea. During the Late Quaternary marine transgression the sea invaded an area of parallel limestone ridges, separated by deep, wide valleys, and created a coast with many bays and long islands parallel and close to the shore. Karstic relief is characteristic of the limestone coast, the outcome of solution processes on rock formations bare

of soils and vegetation. The rock surface has been partly dissolved to leave rugged surfaces, and solution penetrates along joints and fissures to form clefts and caves. In many places along the Dalmatian coast the karstic limestone cliffs plunge into the dazzling blue waters of the Adriatic Sea, but there are pockets of coastal lowland and small valleys where vines and olives are grown.

References Baulig H (1930) Le littoral Dalmate. Ann Géogr 219:305–310 Cvijič J (1924) Geomorfologija. Belgrade De Martonne E (1948) Traité de Géographie Physique [see the chapters on La Côte Dalmate, La Plaine Pannonique (éventuellement le Karst) and La République Fédérative de Yugoslavie]. Armand Colin, Paris Ford D (2007) Jovan Cvijic and the founding of karst geomorphology. Environ Geol 51:675–684 Paskoff R (2005) Karst. In: Schwartz ML (ed) Encyclopedia of Coastal Science: London, Baker & Taylor, pp 581–585 Shepard FP (1973) Submarine Geology. Harper & Row, New York Von Richthofen F (1885) Führer fur Forschungsreisende [see Die ­Dalmatsiche Küste]. Jarecke, Hanover, Berlin

8.19 Croatia

Roger Charlier

The coast of Croatia (Hravatska) extends from the fortified seaport of Pula at the southern tip of Istria (Slovenia) to the Gulf of Kotor (Cattaro) (Montenegro). The Dinaric Alps lie behind a narrow coastal plain, which includes scarp-foot pediments. They follow the geological strike, NW to SE, in folded rock formations, and end in steep slopes overlooking narrow coastal plains, or descending straight to the sea. South of the Istria Peninsula, the Dalmatian coast is bordered by many elongated islands that follow the structural trend of the Dinaric Alps. Narrow peninsulas

protrude into the Adriatic Sea between elongated straits and bays: the channels (canali) zigzag in directions that are mostly parallel, but occasionally at right angles to the coastline (De Martonne 1948; Von Richthofen 1885). During the Late Quaternary marine transgression, the sea invaded an area of parallel limestone ridges, separated by deep, wide valleys, and created a coast with many bays and long islands parallel and close to the shore. Karstic relief is  characteristic of this limestone coast, the outcome of solution processes on rock formations bare of soils and vegetation. The rock surface has been partly dissolved to

⊡⊡ Fig. 8.19.1 Part of the Dalmatian coastline.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.19, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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leave rugged surfaces, and solution penetrates along joints and fissures to form clefts and caves. In many places along the Dalmatian coast, the karstic limestone cliffs plunge into the dazzling blue waters of the Adriatic Sea, but there are pockets of coastal lowland. In general, the Croatian coast is steep and rocky, with  occasional pocket beaches of sand or gravel. The coasts of the Istria Peninsula and the Rijeka (Fiume) ­district on Kvarnes Bay to the east are steep, and penetrated by narrow inlets (rias) along drowned valleys (>Fig. 8.19.1). There has been much development along the coastline. Southeast from Rijeka, the coast is a steep slope (>Fig.  8.19.2) followed by a coastal highway. There are occasional shingle beaches (>Fig.  8.19.3), usually at or near the mouths of steep valleys where streams bring

⊡⊡ Fig. 8.19.2 The steep coast of Croatia north of Senj. (Courtesy Geostudies.)

down gravels when discharge is high after heavy rain. Offshore are the high islands of Krk and Cres, the outer coasts of which are exposed to strong wave action, whereas the mainland coast is very sheltered. To the south are the elongated limestone islands of Rab and Pag, with valleys and inlets running along the NW–SE geological strike. They are dry, with extensive areas bare of vegetation, and the limestone outcrops are rugged. Cliffs on the seaward side show karstic dissection (>Fig. 8.19.4), with solution notches at present sea level, while on the inner shores slopes descend to sea level, incised by valleys that are essentially wadis. The lower slopes and coastal plains have a mantle of loess, locally eroded into low cliffs, which have yielded sandy material for occasional bay beaches. Behind the islands is the Kelebitski Canal, a strait that continues southeast to the mouth of the Zrmanja valley, ending in the almost enclosed Novigrad Lagoon, which has bouldery limestone shores. Some bay mouths have been enclosed by dams to form artificial ponds in which sea water evaporates to brine and crystalline salt (>Fig. 8.19.5). At Povljana, the seaward slopes carry scrubby veg­ etation while the steeper landward slope is a dry escarpment,  with springs issuing from the scarp foot. There are  low cliffs, partly in the loess capping over dipping limestone across which shore platforms have been cut, and locally the beach sand has been cemented to beach rock (>Fig. 8.19.6). The deep Paklenica canyon opens from the high mountain range on to a lobate foreland, which originated as a conical debris fan, now incised by the river that swings southeastward. Low cliffs expose silicified gravels, which locally form a shore bench, and there are gravelly beaches (>Fig. 8.19.7 ). Near the point, the coast has been cut back sufficiently to remove most of a Roman fort. Another valley-mouth lobate foreland at Seline has a watercourse opening on the northern shore, which delivers gravels to a shingle beach that drifts south-east, around the point. Low cliffs to the south expose a conglomerate ce­­mented by calcareous material and a soft chalky limestone. The coast stands forward from the Zrmanja valley, where hill ridges continue the trend of Pag island behind the seaport of Zadar. There are high cliffs on the seaward side of offshore islands such as Kornat. The deep Kirka valley opens into an estuarine inlet, Skradinski Buk, widening down to Sibenik (>Fig.  8.19.8). To the southeast, the coast lacks fringing islands and is exposed to Adriatic storm waves, which have cut shore benches.

Croatia

This is the true Dalmatian coast, extending southeast to the Neretva River. The seaport of Split is sheltered by islands, and to the southeast are the large islands of Brac, Hvar, and Korcula, and many smaller islands. Pocket  beaches of sand and gravel occur locally, and in some places small islands have been tied to the mainland by the formation of tombolos. The coast of Croatia is unique in that it is paralleled by elongated islands but ⊡⊡ Fig. 8.19.3 The coast near Senj, showing pebble beaches. (Courtesy Geostudies.)

⊡⊡ Fig. 8.19.4 Karstic limestone cliffs on the island of Pag, showing dissection along joints. (Courtesy Geostudies.)

8.19

occasionally islands are tombolos. In parts of Brac (Vidova Gove Mountain), there are pine woods, vineyards, and olive groves, while Brac marble is  quarried at Pucisca, Selca, and Sumartin. Numerous bays and beaches are found on Hvar. The island of Korcula, south of Hvar, also has quarries of white marble. The town of Korcula, on the eastern side of the island, has a long history, and there are several ­pebble

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⊡⊡ Fig. 8.19.5 Salt pans in a valleymouth inlet on Pag (Solana Pag). Established in the fifteenth century, these have presumably been adapted to changes in the level of the Adriatic Sea. (Courtesy Geostudies.)

⊡⊡ Fig. 8.19.6 Shore platform cut across dipping limestone at Povljana, showing horizontal layers of beach rock. (Courtesy Geostudies.)

beaches. To the east, the Neretva River has built a delta into a narrow bay behind the long peninsula of Peljesac. The island chains converge on the mainland coast towards Dubrovnik, an ancient seaport, and there are cliffs along the base of the steep coast to the southeast, exposed

to the open Adriatic Sea. Across the border, in Montenegro, the relief is interrupted by a deep curving ria at Boka Kotorska (so-called Gulf or Fjord of Kotor, but resumes and continues to the broad Bojana (Buene) valley, where the Albanian border follows the river down to the coast.

Croatia

8.19

⊡⊡ Fig. 8.19.7 A shingle beach at Starigrad. (Courtesy Geostudies.)

⊡⊡ Fig. 8.19.8 Skradinski Buk, a ria on a limestone coast in Croatia. (Courtesy Geostudies.)

References Baulig H (1930) Le littoral Dalmate. Ann Géogr 219:305–310 De Martonne E (1948) Traité de Géographie Physique [see the chapters on La Côte Dalmate, La Plaine Pannonique (éventuellement le Karst) and La République Fédérative de Yugoslavie]. Armand Colin, Paris

Kotar V (1979) Islands of the Adriatic. Summerdale Press, Milan, Italy Von Richthofen F (1885) Führer fur Forschungsreisende [see Die dalmaticshe Küste]. Jarecke, Hanover, Berlin

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8.20 Bosnia-Herzegovina

Roger Charlier

Bosnia-Herzegovina has a short (20 km) sector of Adriatic coastline, including a segment of the Peljesac peninsula. It is steep, backed by mountains, and shows

karstic features similar to those of Croatia. Poljes dominate the very narrow Adriatic shore.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.20, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.21 Montenegro

Roger Charlier

Montenegro (Cnra Gora) is a mountainous, forested country. Its coastline is only about 48 km long, and includes the deep, steep-sided Gulf of Kotor. It is a generally steep coastline, dissected by many bays (Tivat, Risan, Kotor), linked by natural channels. There are depositional sectors, notably toward the Albanian border,

where the Buna River has supplied sediment that has drifted northward. The Adriatic coast here has a microtidal regime (tide ranges less than 50 cm). Longshore currents usually are relatively slow, ranging between 30 and 200 cm/s.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.21, © Springer Science+Business Media B.V. 2010 (Dordrecht)

8.22 Albania

Yurii Shuisky

1.  Introduction

⊡⊡ Fig. 8.22.1 Coastal geomorphology of Albania.

The coast of Albania, from the mouth of the Buna (Buene) River south to Cape Stylos in Kerkira Strait is about 385 km long. There are steep sectors of coastline, mainly along the seaward flanks of ridges that trend NW–SE, and there are low-lying sectors, chiefly at the mouths of river valleys where deltas have been built. The ridges end in capes, many of which are in encephalon sequence, sheltering bays that open northward (>Fig. 8.22.1). The continental shelf narrows southward along the Albanian coast, and almost disappears along the Karaburuni Peninsula, where the relatively short westerly fetch across the Adriatic Sea gives place to a longer southwest fetch across the Ionian Sea. Water temperatures range from about 12°C in February to 25°C in August. The mean salinity of the sea is 3.5%, decreasing near river mouths: it is lower in winter, when river discharge is larger, and higher in summer. Tide range is very small, less than 10 cm, but strong winds from the southwest raise water levels and generate waves 2–3 m high from that direction on the outer coast. It has been shown that wave action is effective in moving sea floor sediment to a depth of 8 m off the Albanian coast. Waves have cut cliffs on headlands but there has been progradation in many of the bays as the result of deposition of sand and gravel brought down by rivers and eroded from nearby cliffs (Paskoff 1985). Cliff recession rates of up to 0.3 m/year have been measured. It has been estimated that during the past century, 28% of the coastline has prograded, 42% has been cut back by erosion, and 30% has remained more or less stable. Cliff erosion has yielded an average of 228,000 m3/year, while the combined yield of Albanian rivers is 52.9 million m3/year, of which 25% is relatively coarse material retained on beaches and 75% silt and clay dispersed offshore. Progradation has resulted in the formation of a wide dune-fringed coastal plain (Doody 2005) and the stranding of the town of Lesha, which stood on the coast in the first century ad and is now 8 km inland behind Drin Bay. The ancient town of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.22, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Apollonia stood 5 km inland when it was founded 5,000 years ago, and is now 8 km inland.

2.  The Albanian Coastline The Buna delta on the Yugoslav border is cuspate and south-facing, with delta-flank creeks and lagoons on either side. From the mouth of the Buna River, the lowlying coast curves south east, along the southern side of a  coastal ridge and round to the large Drin delta. There are many beach ridges indicating stages in progradation, and a large lagoon on the southern side of the Drin delta. To the south is another delta, that of the Matya River, before the coast runs out to Cape Rodoni. The projecting  headlands are on Cretaceous and Tertiary rock formations. The bay of Lalzes then curves southward to the Erzen delta, and out to Cape Palit, a narrow protruding ridge. Then a sector of hilly coast runs south to Durres, where a curving breakwater shelters a harbour. Durres Bay extends round to Cape Lagit (Lagy), where another hilly ridge projects into the sea. The Shkumbo River flows to the sea on a west-facing coast, and has built only a blunt delta. To the south, a barrier spit encloses the Karavastas Lagoon and the Seman River flows to a delta on the southern side. There are beach

ridge plains with some dunes. Broad arcuate bays run out to the cuspate remains of abandoned deltas on the coastal plain, then the Vijose delta, with a beach ridge plain to the south. This is interrupted by the Nartes Lagoon, behind spits that link former small high islands. The Gulf of Vijose occupies a trough behind the high Karaburuni Peninsula, which runs out NNW to Cape Gjuhezes, with the high outlying island of Sazan. The seaward slopes of the Karabuni Peninsula are steep, and the nearshore sea deep, with exposure to strong southwesterly waves across the Ionian Sea. Cliffs up to 50 m high have been cut in the strongly folded Cretaceous limestones and sandstones. Karst topography has developed locally on limestone shores. The steep coast follows the trend of the mountain range southeast to Cape Kefali, a cliffed promontory beyond which the coast passes into the shelter of the island of Corfu. Lake Butrinut lies within a deltaic plain close to the Greek border at Cape Stylos.

References Doody JP (2005) Europe, coastal ecology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 442–452 Paskoff R (1985) Geomorphological aspects of the coast of Albania. (in French) Bulletin, Association de Géographes Français 2:77–83

8.23 Greece

1.  Introduction Greece has a 13,676 km coastline, which is long for its area (131,944 sq. km). The country has been, and still is, strongly affected by tectonic and volcanic activity. The Greek peninsula was the outcome of the Alpine orogeny: it was previously part of the Tethys Sea, but with the Alpine orogenic uplift there was folding and faulting along the NW–SE trend of the Hellenides (Kronberg and Günther 1978), which led to an extremely rugged topography, with a large number of horsts and grabens (fault-bounded ­corridors of uplift and subsidence, the horsts forming promontories and the grabens bays), extending NW–SE. Vertical displacement between the crests of these horsts and the floors of the grabens can be up to 2 km. In addition, major upwarping of the Greek peninsula has led to  high spine-like mountain ranges such as the Pindos Mountains inland and the Taygetos Mountains, which run out as a peninsula to Cape Matapan (Akrotirion Tainaron) in the south-­central Peloponnese. There has also been E–W faulting, leading to horsts and grabens in that direction: for example, the Gulf of Corinth, which separates the Peloponnese from the mainland, and the Gulf of Malia on the east coast of Greece. The Aegean Sea is the locus of the collision of several continental plates, and is thus a very complex tectonic setting. Tectonic movements have continued, displacing former coastline features upward or downward around the coast (Lagios and Wyss 1983). The coastal climate of Greece is Mediterranean, with mild rainy winters and warm dry summers. Rainfall is higher in the west, where Corfu has 1,320 mm. Athens has a mean monthly temperature of 8.6°C in January, rising to 28.2°C in July, and an average annual rainfall of 414 mm. Tide ranges are small, less than 1 m. The predominant waves that affect the coasts of Greece are produced by NNE (Meltemia) winds, but storm-generated waves also arrive from the south and west over the Mediterranean Sea. Much of the coastline of Greece is high and steep, with cliffs of varying dimensions. Where the cliffs are cut in

hard limestones or metamorphic rocks, erosion has proceeded very slowly, but where they are cut in soft Neogene sands, silts, and marls, erosion can be extremely rapid. Cliffing is also related to exposure, being more prominent where the coast receives strong wave action. There are also many sedimentary embayments, some with prograding deltas. Much of the land surface is un­­ dergoing intensive erosion, which accelerated during the Holocene Epoch and in particular the past 5,000 years, following deforestation and grazing. Consequently, large quantities of sediment have been delivered to the coast by runoff and rivers. Sediment have flowed into the grabens, filling from the landward ends of these deep tectonic  features, but the grabens have not been filled the completely, as massive hinterland erosion and fluvial delivery of sediment have occurred only in the past few 1,000 years. Greek terms for coastal features include akrotirion (cape), dhiavos (strait), khersonisos (peninsula), limni (lake), nisos (island), ormos (bay) and potamos (river).

2.  The Coastline of Greece South from the Albanian border the Ionian Sea coast of Greece is generally steep, with low-lying deltaic areas at valley mouths. The Thiamis River has built a delta into the sheltered waters of the Strait of Kerkir in the lee of the mountainous island of Corfu. On the west coast of Corfu and the smaller islands (Paxoi and Andipaxoi) to the south the steep slopes show  basal cliffing with exposure to southwesterly wave action over the Ionian Sea (>Fig. 8.23.1). The steep coast extends south of Parga, interrupted by the Akheron delta which has formed behind a breached coastal ridge, and at Preveza a narrow strait leads into the large branching Amvrakikos Kolpos, a landlocked bay. On its northern shore is the Louros delta, with curved sandy spits and barriers enclosing lagoons and marshes to the east, and the Arakhthos delta, while the southern coast has steep-sided

Edited version of a chapter by C. Tziavos and J.C. Kraft in The World’s Coastline (1985: 445–453). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.23, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.23.1 Limestone cliff on the west coast of Corfu. (Courtesy Geostudies.)

promontories and narrow bays. South of Preveza a depositional lowland contains a lagoon, Pounda Salton, and a canal links swampy Lake Voulkaria to the sea. To the south the bold west coast resumes, with steepsided promontories and small bays mainly at valley mouths. The high island of Levkas and the similar islands of Cephalonia and Zante to the south each have steep limestone coasts and some bay beaches. On the mainland coast the Akarnanika Range runs southward to a high promontory, backed by a deep valley opening to the sea at Astakos. An intricate ria coast fringed by small islands south of Astakos gives place to the large Akheloos delta with its fringing barrier islands and lagoons beside the Patraikos Gulf. A bay with an intricate shoreline extends eastward to the Evinos delta as the Gulf narrows into the strait at Andirrion, which then widens into the deep, steep-sided Gulf of Corinth. Shingle beaches south of the Corinth isthmus have fragments of antique brown pottery mixed with white limestone pebbles. East of the Mornos delta the north coast of the Gulf of Corinth is irregular, with high promontories and deep valley-mouth gulfs, whereas the south coast is relatively straight, with a narrow, sometimes lobate, coastal plain (Soter 1998). This fringing lowland continues past Patras on the south coast of the Patraikos Gulf and out to Cape Papas, where a ridge borders a narrow inlet. The coast then swings southward in a gentle curve with small headlands at as Bania Kounopeli and beaches which become barrier spits on either side of a marshy lagoon south of Manolas. The coast curves out to a cliffed promontory of

Pliocene sedimentary rocks at Kyllini, extending round to  Tripiti and facing across Zakinthos Strait to the high island of Zante. The next bay faces southwestward, and has sandy beaches fronting a beach ridge plain on either side of the Pinios River. Cliffs resume as the coast swings southward to another hilly promontory of Pliocene rocks at Katakolo, where the receding cliffs are exposed to strong southwesterly wave action, the fetch across the Medi­ terranean Sea to Tripoli being 900 km. To the south is the wide Kiparisiakos Bay, backed by long, steep sandy beaches on either side of the blunt delta of the Alphios River (Poulos et al. 2002). The beach is backed by beach ridges indicating progradation (Raphael 1969), and stands as a barrier in front of the lagoons of Muria west of the delta and Agulinitsa and Kaiafa to the east, all of which have extensive marshes on their inner shores. The coastal plain narrows southward, continuing past Kiparissia and behind Proti Island on the west coast of the Messenian peninsula to Oxbelly Bay and the Eocene limestone upland of Paleokastro. Parts of the coastline include Pliocene deposits. The beaches are narrow, and the deep waters of the Ionian Sea lie close to the shore. To the south the elongated high island of Sphakteria, another ridge of Eocene limestone, shelters Navarino Bay. Large quantities of sand drifting along the Ionian Sea coast have accumulated to seal off a former channel into the present Osmanaga Lagoon with a barrier bearing an arcuate sandy dune. The Sikia channel between Paleokastro and Sphakteria, was deep at the time of the Battle of Lepanto (1571) between the Turks and the Venetians, but a line of ships was sunk in

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the channel to stop access to Navarino Bay shortly after this battle, and this formed a baffle that caused sand to accumulate and reduced the channel to a shallow depth. The southern entrance to Navarino Bay remains deep, but is narrow, limiting wave action from the Ionian Sea. Navarino Bay, which is probably a small graben, has soft Pliocene sands and silts on its eastern shore. A number of rivers flow into Navarino Bay, and fluvially-supplied sediment is transported by longshore drift set up by waves largely generated within the almost enclosed em­­bayment. There is evidence that Navarino Bay was larger about 8,000 years ago, when it extended at least 4 km ­farther north than at present (Kraft et al. 1980). A broad sandy barrier spit has grown across the northern part, separating Osmanaga Lagoon (>Fig.  8.23.2), which is bordered by extensive marshes to the east, an alluvial plain to the north and a dune-covered sand barrier behind Oxbelly Bay to the west. Archaeological evidence indicates that this spit already existed 2,000 years ago. Osmanaga Lagoon (>Fig.  8.23.2) is subject to strong evaporation, and its muddy sediment are at least 50% calcium carbonate. The sandy terrain to the northwest of Navarino Bay includes large amounts of beach rock (containing Roman shards) and older Quaternary sands in an area that has been tectonically raised. A large amount of sand has moved inland in the form of coastal dunes. South of Navarino Bay a steep coast of hard Eocene limestone extends along the western side of the Messenian peninsula. In the bay at Methoni soft Neogene sediment outcrop in rapidly receding cliffs (>Fig. 8.23.3). To the east the south coast of the Messenian peninsula consists of Triassic siliceous shales and Eocene to Miocene shales and conglomerates, while harder Cretaceous rocks outcrop on the western side of the Gulf of Messenia from Coroni north to Petalidi. The Messenian Gulf is another example of the filling of the head of a major graben. There is exceptionally high relief between the deep graben of the Messenian Gulf and the high mountain range, the Taygetos, to the east. There is erosion on beaches in the northwest of the Gulf. The coast at the head of the Gulf of Messenia, from Petalidi to Kalamata in the east, is a curving sand and gravel beach receiving

⊡⊡ Fig. 8.23.2 The lagoon at Osmanaga north of Navarino Bay, with the deep water embayment of Navarino facing Eocene limestone ridge of Paleokastro.

8.23

sediment from the large Pamisos River and five smaller rivers incised into Neogene sediment on the western side of the head of the Gulf. The eastern coast of the Messenian Gulf consists of high cliffs and steep slopes, mainly of hard Cretaceous limestone, with some very small cove beaches, and to the south there are numerous small rias. The Taygetos Range ends at Cape Matapan, and on its eastern side steep embayed coasts run alongside the La­­ konian Gulf. This is similar to the Messenian Gulf in having a depositional coastal plain with beaches and dunes at its head, on either side of the Evrotas River (Cundy et al. 2006), and a high peninsula on the eastern side, running south to Cape Zovello. The coast then turns northward along the Mirtoan Sea, and is generally steep with minor cliffing and some valley-mouth lowlands. An indented steep coastline runs along the western side of the Argolis Gulf, which also has a depositional coastal plain at its head and an embayed eastern coast with the outlying islands of Spetsai and Idhra. Beyond this peninsula the coast runs northwest past Poros and the narrow-necked Methana peninsula along the embayed steep coast of the Saronic Gulf, composed of Mesozoic limestones, with small pocket beaches of sand. There are no major rivers in this part of Greece. The high country recedes, and the 6 km Corinth Canal has been cut through a narrow isthmus to the Gulf of Corinth, where there has been tectonic deformation (Mariolakos and Stiros 1987). East of the Corinth Canal the north coast of the Saronic Gulf is low-lying, with small beaches between low cliffs of Pliocene and Quaternary age. This is followed by very steep cliffs of Cretaceous and Triassic limestones east to Mergara, then rocky promontories of Mesozoic shale and marble and Neogene sedimentary rocks separated by small bays with curving beaches. A coastal lowland widens in the lee of the island of Salamis, and the coast becomes urban and industrialised between here and Paleo-Phaleron, southeast of Athens. The largest rivers are the Kiphisos and the Ilissos, which have formed major alluviation of the plain now completely covered by the city of Athens. There is little sand along the east coast of the Saronic Gulf, and as this is a major resort area there has been

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⊡⊡ Fig. 8.23.3 Cliffs on the east coast of the bay at Methoni in the SW Peloponnese.

a­ rtificial beach nourishment. The coastline east of Athens is one with major formations of beach rock. Attica lies in a semiarid zone but with sufficient rainfall to support a vegetative cover. Unfortunately, human activity has destroyed much of this cover, and massive erosion has occurred over the past 5,000 years. There are extensive artificial salt-making lagoons behind Anavissou Bay. The hilly coast curves round Cape Sounio and runs northward beside the strait that separates Makronisos Island, a long and narrow ridge, trending SSW–NNE and rising to 264 m. To the north are valley-mouth lowlands as far as Marathon Bay, which faces SE and ends in a narrow projecting ridge out to Cape Marathon. Strong northward drifting has supplied sand to a beach ridge plain that stands seaward of the marshes of the Marathon Plain, famous as the trap into which the Persian army was manoeuvered and defeated in 490 bc. Although the coastal processes at Marathon seem very active, the coastal area is still much as it was described 2,500 years ago (Kraft 1972). A narrow coastal plain runs along the southwestern side of the Notios Evvoikos Gulf, past the Asopos delta to the narrow winding strait at Kalkis, and similar low sectors alternate with high and steep coasts around Evvoia. The strait has the so-called euripos tides, which produce strong and variable current flow. At the northwest end a narrow strait leads in to landlocked Vorios Evvoikos Gulf, which again has steep coasts alternating with coastal plains, particularly at the mouths of valleys (Pirazzoli et al. 1999). Atalanti Bay was the site of earthquake subsidence in 1894,

causing slumping, tsunami flooding and the separation of the Gaidurosini peninsula from the mainland: elucidated by wetland biostratigraphical studies (Cundy et al. 2000). A western arm is Maliakos Gulf, which was much larger in 480 bc, when the Battle of Thermopylae, between Greeks and invading Persians, was fought along a narrow coastal plain (Pass of Thermopylae) below a steep slope rising to Kallidromon Mountain on the southern side. Subsequently sediment deposited by the Sperchios River has built a large deltaic plain in front of the Pass of Thermopylae. Landlocked Pagasitikos Bay is bordered by irregular peninsulas and inlets, with coastal plains north of Almiros on the west coast and around the port of Volos in the north. The Mavrovouni Range runs southeast through the peninsula on the eastern side, and a chain of high islands, the Northern Sporades (including Skiathos, Skopelos, Iliodhromia and Pelagos) curves out into the Aegean Sea. A mainly steep coast with low cliffs cut in Pliocene rocks and an intermittent narrow lowland fringe runs northwest alongside the mountain ranges, which are breached by a deep gap north of Agiokambos and the Timbios gorge, which opens to a delta at Cape Platamon. Behind Litokhoron the land rises gradually, then steeply to Mount Olympus (2,911 m), the highest mountain in Greece. The coastal plain then broadens towards the head of Thermaikos Gulf, with a lobate foreland (Akrotirion Atheridha) near Katerini, formed as a looped spit with recurves into a lagoon on the inner shore. Thermaikos Gulf is a large graben. Several large rivers,  including the Axios, enter the embayment from the

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north,  and their large load of sediment has filled hundreds  of square kilometres of the former head of the Thermaikos Gulf. The ancient capital of Macedonian Greece, Pella, was a port in the time of Philip and Alexander, but now lies about 40 km inland. Growth of the Aliakmon and Axios deltas resulted in their convergence, separating part of the former Thermaikos Gulf from the sea as Lake Ludias. from the rapidly prograding coast of the Axios delta. This lake has become a largely reclaimed marsh area. It is likely that further delta growth will cut off more of the Thermaikos Gulf as a lagoon near Salonika within the next 50–200 years, unless this is prevented by human actions. Fluvial sands carried across the sea floor and along the coast have formed extensive beaches lie in the area of Aghia Trias and Peraia and a cuspate foreland at Epanomi. East of Salonika the Khalkidhiki peninsula has three steep-sided ridged prongs, Kassandra, Sithonia and Akti,

projecting into the Aegean Sea and separated by Toronaios and Singitikos bays. Steep coasts predominate here and to the east, interrupted by the valley of the Strimon River. The large Nestos delta has been built in the shelter of the high round island of Thasos, and to the east Lake Vistonis has been impounded by a spit in a former embayment. The coastal plain of Thrace is diversified by hilly ridges, extending to the Maritsa (Meric) delta, where the river forms the boundary with Turkey.

3.  The Greek Islands The Greek islands of the Aegean Sea are highly varied in their geology, their coasts being mainly rocky cliffs with small sandy beaches. Some of are limestone, others metamorphic or volcanic rock.

THERA ISL. (SANTORINI)

Contour interval 100 Metres Sandy Shorelines Cliffs Prevolcanic

L.

Mostly Triassic Limestone

ER

AS

SI

A

IS

Active Volcano

TH

TH

⊡⊡ Fig. 8.23.4 The Island of Santorini, showing submarine contours that remained after the massive explosion of the caldera about 3,000 years ago and the new volcanic cone developing in the centre. (Courtesy Geostudies.)

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ER

m

G AE

C

N IA

LD

18

A SE

18

A

E

R

A

NE

A

CA

M

E

IS NI

L.

0m

MESA VOUNO

0m

0

2

N

A

360

4 Km

ANCIE

NT TH ERA

779

780

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⊡⊡ Fig. 8.23.5 The volcanic cliff at Thira, Santorini. (Courtesy Geostudies.)

⊡⊡ Fig. 8.23.6 The cliff at Thira, viewed from the active volcanic cone in the centre of Santorini. (Courtesy Geostudies.)

To the north lies the Island of Santorini or Thera or Thira (>Fig. 8.23.4), a volcanic island formed of a composite cone that exploded about 3,200 years ago with a force estimated to have been tens of times as great as that of Krakatau in Indonesia. As at Krakatau, remnants of the original island surround a crater invaded by the sea. Some historians and archaeologists claim that the resultant tsunamis wiped out the Minoan civilisation on Crete. In any case, the very heavy fall of pumice and tephra deeply covered the landscape of the island and buried an

important city on the southeast of the island. Thera now comprises only part of the larger, former volcanic cone, but submarine contours show the shape of the cone and the very deep central caldera, and there are steep infacing cliffs, as below the town of Thira (>Fig.  8.23.5). Other small islands form remnants of the outer edge of the volcanic cone, and within the caldera lies Nea Kamini, a new volcanic cone that has been intermittently active and growing within the past 300 years, with lava flows down to the sea (>Fig. 8.23.6).

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⊡⊡ Fig. 8.23.7 Steep cliffs in basaltic lava on the south coast of Santorini. (Courtesy Geostudies.)

The outer coasts of Santorini have generally low cliffs cut in tuff (volcanic ash), some of which is indurated with beaches of grey volcanic sand and gravel, mixed with white calcareous sand. Locally there are bolder cliffs in basaltic lava (>Fig. 8.23.7). Northwest of Crete is the small island of Antikythira, which has a limestone cliff showing a set of nine emerged shore horizons between 1.1 and 2.7 m above sea level, indicating a series of small subsidence events followed by  a sharp uplift during a major earthquake in 365 ad (Pirazzoli et al. 1982). Further north are the Cyclades, an archipelago of high islands (the peaks of a largely submerged mountain range) including Amorgos, where Stiros et al. (1994) studied the effects of tectonic movements. To the northeast the Dode­ canese archipelago includes the large island of Rhodes, where tectonic movements of Holocene shorelines were discussed by Pirazzoli et al. (1989). There are multiple bioerosion notches at several levels up to 2.5 m, indicating intermittent uplift. The coast is mainly rocky, with some bays and marine inlets, and there are beaches of generally local provenance (Pyokari 1997). To the northwest the island of Nisiros is an extinct volcanic cone and to the southwest is the elongated high island of Scarpanto. In the northern Aegean Sea are more high islands, Samos, where emerged Holocene shore terraces have been studies (Stiros et al. 2000), Rhios and Lesvos, close to the Turkish coast, and Skiros, Evstratios, Lemnos and the volcanic island of Samothrace in line SSW–NNE.

References Cundy AB, Korekaas S, Dewez T et al. (2000) Coastal wetlands as ­recorders  of earthquake subsidence in the Aegean: a case study of the  1894 Gulf of Atalanti earthquakes, central Greece. Mar Geol 170:3–26 Cundy AB, Sprague D, Hopkinson L et al (2006) Geochemical and stratigrapohic indicators of late Holocene coastal evolution in the Gythio area, southern Peloponnese. Mar Geol 230:161–177 Kraft JC (1972) A reconnaissance of the geology of the sandy coastal areas of eastern Greece and the Peloponnese. Technical Report 9, College of Marine Studies, University of Delaware Kraft JC, Rapp GR, Aschenbrenner SE (1980) Late Holocene paleogeomorphic reconstructions in the area of the Bay of Navarino: Sandy Pylos. J Archaeol Sci 7:187–210 Kronberg P, Günther R (1978) Crustal fracture pattern of the Aegean region. In: Close H, Roeder D, Schmidt K (eds) Alps Apennines Hellenides, Inter-Union Commission on Geodynamics scientific report 38. E. Schweizerbart’sche Verlagsbuch­handlung Stuttgart, Germany, pp 522–526 Lagios E, Wyss M (1983) Estimates of vertical crustal movements along the coast of Greece, based on mean sea level data. Pure Appl Geophys 121:869–887 Mariolakos I, Stiros SC (1987) Quaternary deformation of the Isthmus and Gulf of Corinth. Geology 15:225–228 Pirazzoli PA, Laborel J, Laborel-Deguen F (1999) Late Holocene ­coseismic vertical displacements and tsunami deposits near Kynos, Gulf of Euboea. Central Greece. Phys Chem Earth PT A 24:361–367 Pirazzoli PA, Montaggioni LF, Saliège JF, Segonzac G, Thommeret Y, Vergnaud-Grazzini C (1989) Crustal block movements from Holocene shorelines: Rhodes Island, Greece. Tectonophysics 170:89–114 Pirazzoli PA, Thommeret J, Thommeret Y, Laborel J, Montaggioni LF (1982) Crustal block movements from Holocene shorelines: Crete and Antikythira (Greece). Tectonophysics 86:27–43

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Poulos G, Evans G, Voulgaris G (2002) Sediment flumes and the evolution of a riverine-supplied tectonically-active coastal system: Kyparis­ siakos Gulf, Ionian Sea. Geological Society (Special Publication) 191:247–266 Pyokari M (1997) The provenance of beach sediments on Rhodes, southeastern Greece, indicated by sediment texture, composition and roundness. Geomorphology 18:315–332 Raphael CN (1969) The Plain of Elis, Greece – an archaeological approach. Mich Acad Sci 1:73–74

Soter S (1998) Holocene uplift and subsidence of the Helike delta, Gulf of Corinth, Greece. Geological Society (Special Publication) 146:41–56 Stiros SC, Marangou L, Arnold M (1994) Quaternary uplift and tilting of Amorgos Island (southern Aegean) and the 1956 earthquake. Earth Planet Sci Lett 128:65–76 Stiros SC, Papageorgiou S, Evin J (2000) Seismic coastal uplift in a region of subsidence: Holocene raised shorelines of Samos Island, Aegean Sea. Mar Geol 170:41–58

8.24 Crete

Anja Scheffers · Tony Browne

1.  Introduction Crete is the largest of the Greek Islands, with an area of 8,259 km², about 260 km from west to east and ranging between 12 and 60 km north to south. It is ­situated in the southern Aegean Sea, just north of latitude 35° N. The coastline is about 1,000 km long. Together with the neighbouring islands of Kithira and Antikithira to the NW and Kos, Karpathos and Rhodes to the NE, Crete forms an island arc (the Aegean Arc), parallel to the Hellenic Trench in the subduction zone formed by the collision between the European and African plates. It is tectonically active, and was much affected by earthquakes between the mid-fourth and mid-sixth centuries ad (Pirazzoli et al. 1996). The island is mountainous, rising to about 2,400 m above sea level, with some coastal plains, particularly in the north. Between the mountain ranges are tectonic basins and grabens (Fytrolakis 1980). Limestones ranging from early Palaeozoic to mid Jurassic age are extensive and karst features such as polje and caves are common. Around the perimeter of the island, Neogene carbonate sandstones occupy tectonic basins, and in the Pleistocene there was extensive faulting, with areas of uplift in the east, south, and south-west and subsidence along much of the north coast. There has been continuing tectonic deformation. Coastal dunes and dune calcarenites extend behind beaches on the north coast. The climate is Mediterranean, with hot dry summers and wet, mild winters, particularly along the western slopes of the mountains. Runoff is augmented by the melting of winter snow in spring, but river discharge is meagre in summer. Prevailing winds are westerly in winter and alternate between north and south in summer. Spring tide range is usually less than 0.20 m. Wave energy is relatively strong on the west and south coasts, facing the open Mediterranean Sea.

produced largely by bioerosion. Beaches are rare and restricted to bays that open to the north and to coves along the south and west coasts. The majority of beaches consist mainly of pebbles and cobbles. Many beaches have eroded beach rock outcrops. Small biogenic formations, formed by calcareous algae and vermetids, are common on shores in the upper subtidal area (Kelletat and Zimmermann 1991). Dislocated by neotectonics (Peterek et al. 2003), their presence along with bio-erosive notches has been used to determine the date and scale of past earthquakes: for example, the sudden uplift of up to 9 m on the west coast on 21 July 365 ad (Kelletat 1996; Pirazzoli et al. 1996; Scheffers 2006; Stiros 2001). There is evidence of differential uplift at many sites around Crete (Peterek et al. 2003), and also for successive tsunamis, notably the one that occurred about 3,500 years ago (Bruins et al. 2008; Scheffers and Scheffers 2007). Ancient harbours in western Crete have been raised above sea level while harbours and towns of similar age in eastern Crete have been submerged by the sea. The calcareous alga Neogoniolithon notarsii occupies a narrow sublittoral horizon and is a good indicator of past and present sea level (Pirazzoli et al. 1996). At Tripitis on the south coast, there are caves that were raised by the seismic movements in 365 ad, with floors related to the earlier higher sea level. On the NW coast, parallel ridges project into the sea on either side of Kisamou Bay. The north coast is generally steep, with small coastal plains as at Iraklion and several large bays. The NE point, Cape Sidheros, is a narrow winding peninsula. Much of the south coast is steep with some basal cliffing, and segments of narrow coastal plain that widen at Timbakion, the mouth of the Yeropotamus valley (>Figs. 8.24.1 and > 8.24.2).

References 2.  The Coastline of Crete In general, the coastline of Crete is steep and rocky. Cliffs cut in limestone have basal notches and shore rock pools

Bruins HJ, MacGillivray JA, Synolakis CE, Benjamini C, Keller J, Kisch HJ, Kluegel A, van der Plicht J (2008) Geoarchaeological tsunami deposits at Palaikastro (Crete) and the Late Minoan eruption of Santorini. J Archaeol Sci 35:191–212

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.24, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Crete

⊡⊡ Fig. 8.24.1 Uplifted notch with basal double rim of calcareous algae at 7 m above sea level near Falasarna on the west coast of Crete.

⊡⊡ Fig. 8.24.2 Bioconstructional platform built by calcareous algae and vermetids after the uplift of 365 ad along the west coast of Crete.

Fytrolakis N (1980) The geology of Crete island. National Technical University of Athens, Athens, Greece, 80pp Kelletat D (1996) Perspectives in coastal geomorphology of Western Crete, Greece. Z Geomorphol, Suppl.Bd. 102:1–19 Kelletat D, Zimmermann L (1991) Distribution and morphological types of recent and sub-recent organic rocks along the coastlines of  Crete (in German). Essener Geographische Arbeiten 23:168pp Peterek A, Beneke K, Schwarze J, Spinn A (2003) Coastlines and coastal forms in western Crete, reflected by eustatic, tectonic or gravitativetectonic processes (in German). Essener Geographische Arbeiten 35:39–56

Pirazzoli PA, Laborel J, Stiros SC (1996) Coastal indicators of rapid uplift and subsidence. Examples from Crete and other Mediterranean sites. Z Geomorphol, Suppl.Bd. 102:21–35 Scheffers A (2006) Coastal transformation during and after a sudden neotectonic uplift in western Crete (Greece). Ann Geomorphol, Suppl. Bd. 146:97–124 Scheffers A, Scheffers S (2007) First significant tsunami evidence from historical times in western Crete (Greece). Earth Planet Lett 259:613–624 Stiros SC (2001) The AD 365 Crete earthquake and possible seismic ­clustering during the fourth to sixth centuries AD in the Eastern Mediterranean: a review of historical and archaeological data. J Struct Geol 23:545–562

8.25 Bulgaria

Eric Bird

1.  Introduction The Bulgarian coast is 378 km long. It has a complex geological structure, truncating the Balkan mountain ranges (>Fig. 8.25.1). In the south, volcanic formations are overlain by limestones and sandstones in the Strandza anticlinorium, in the central part are highly folded flysch rocks of Upper Cretaceous to early Tertiary age, and in the north the Moesian Platform of almost horizontal Miocene and Quaternary strata, including marls, sandstones, clays, and limestones, with a loess capping (Popov and Mishev 1974). Quaternary marine terraces occur 90–100, 55–60, 35–40, 20–25, 12–14, 4–5, and 1.5–2 m above present sea level, the last two being of Holocene age, the others Pleistocene. There are also submerged Pleistocene marine terraces offshore, beneath the Black Sea. Much of the coastline consists of eroding cliffs, and landslides are extensive, often forming a stepped topography. The stability of coastal slopes varies with geology, limestone forming bold cliffs while soft sediment, including loess, collapses or slumps. Earthquakes may trigger landslides. The impact of human activities is evident in ports and seaside resorts, coastal and undersea quarries, and coast protection works. Sectors of deposition include deltas and beach ridges, and there are several coastal lagoons. In the south, there is a cliffed coast dissected along joint planes (>Fig. 8.25.2). Concave bays are occupied by beaches backed by dunes, especially south of Burgas and north of Varna, where sand from the rivers has nourished prograding beaches and dunes behind indented bays. The intervening steep coast has only pocket beaches. Dunes are extensive in the north, where they are up to 18.6 m high on the 11.2 km long Kamcia–Skorpilov coast. At Nessebar, engineering works have modified the beach and dune system. Sedimentological analyses have shown that beach and dune sands have come mainly from rivers, notably the Batova, Provadiiska, Kamcia, Dvoiniza, Hadziiska, Ruso­ kastro, Ropotamo, Djavolska, Karaagac, Oreska, and Veleka. Grain size diminishes offshore, and landslides have acted

as  breakwaters separating beaches of varying coarseness (Zvetkova and Simeonova 1984). The tide range is negligible, but sea level oscillations accompany changes of wind direction and barometric pressure. Strong onshore winds raise sea level temporarily. Wind regimes on the Bulgarian produce a predom­inance of northeasterly wave action, stronger in the winter months, and generate longshore drifting to the south. However, there has been northward drifting along the Kamcia–Skorpilov beach in northern Bulgaria (Dacev and Nikolov 1977). Severe storms occur from time to time, and the period September 1976 to April 1977 was unusually stormy. Vasilev (1978) recorded the rapid coastline changes that resulted from these storms: beaches were eroded and basal cliffs cut into landslide lobes. There was damage to coastal structures built at Balcik, Varna, Nessebar, and Pomorie, and it seems that human interference made the erosion more severe than it would naturally have been. It is hoped that future engineering works to protect seaside resorts on this part of the Black Sea coast will be based on a better understanding of coastal processes.

2.  The Bulgarian Coast From the Turkish border on the Rezovska Reka, the coast to the north is predominantly cliffed. The cliffs are cut in stratified sandstones, limestones, and clays overlying lava and tuffs, and there has been slumping in the soft capping deposits (>Fig. 8.25.1A). At Ahtopol, the cliffs are interrupted by the valley of the Veleka River. From Micurin northwest to Primorsko and Sozopol the coast is more indented, with rectilinear bays and headlands. The Strandza anticlinorium ends along the southern side of the Bay of Burgas. Sandy beaches and barriers extend past coastal lagoons, Jaz Mandra and Burgasko Ezero, which have outlet channels to the coast. North of Burgas coastal slopes in Mio-Pliocene and Quaternary clays show extensive and rapid slumping (>Fig. 8.25.1D). The Sarafovo landslide forms a lobethat has been cut back by marine erosion, removing volumes of more than 20 m3/m/year. An understanding of the hydrogeology, and

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.25, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 8.25.1 Geology and structure of the coast of Bulgaria (A) – simple landslide on a cliff of loess, (B) – complex landslide at Balcik, (C) – schematic section through a cirque-type landslide north of Varna, (D) – schematic section through the landslide near Sarafovo, formed in a stratified and unstable formation, (E) – schematic section through the landslide north of Micurin. (Courtesy Geostudies.)

especially the pattern and changes in the level of the underground water table is essential for the application of successful engineering works along this slumping coastline (Stanev and Simeonova 1979). The total volume of material lost by ­erosion is some 1,344,000 m3 annually (Shuisky and Simeonova 1982). At Pomorie, there is a sandy tombolo, but the curving beaches alongside it are rapidly eroding. The beach on the eastern side becomes a barrier fronting the Pomorijske Ezero lagoon. Northward, the coast curves north past

Pavda to Nessebar, where the Hazdijska Reka flows out, and then east to Cape Emine. The cliffs at Cape Emine are relatively stable, receding only a few centimeters a year (Simeonova 1985), but on the more complicated Balkan structures to the north, past Obzor and Byma, cliff retreat is as much as 15 m/year. Beyond the mouth of the Kamchiya River valley (where there is no delta), the cliffs are much disturbed by slumping. They decline to Galata, where the strongly folded Balkan structures come to an end.

Bulgaria  

⊡⊡ Fig. 8.25.2 A gulch cut along parallel joints in limestone on the cliffs of southern Bulgaria. (Courtesy G. Simeonova.)

   8.25

the limestone, which caps the plateau (>Fig. 8.25.1C). The stratified cliffs show the effects of weathering and undercutting, leading to recurrent sliding at rates of up to 1 m/ year. The coast curving eastward past Balcik to Kavarna also has extensive landslides in calcareous marls and clays (>Fig. 8.25.1D). The cliffs are up to 20 m high and bolder where there is a capping of limestone, as on Cape Kaliakra, beyond which the coast runs northeastward. North of Cape Kaliakra, the large Tauk Liman landslide is of block slide type, and shows seasonal variations in mobility, but average cliff retreat has been only 0.01 m/ year. Near Cape Sabla, there have been rockfalls on the limestone cliffs, which have retreated at an average of 0.01 m/year. Erosion is more rapid where the sea reaches overlying loess, which breaks away and collapses on to the limestone shore (>Fig. 8.25.1E). North of Cape Sabla, erosion is also more rapid where clays underlying the limestone are exposed, and reaches rates of 8 m/year locally. The slumping cliffs continue to the Romanian border.

References

Varna stands on the northern shore of the Varnenski estuary, and marks the beginning of dissected hilly country on the almost horizontal Miocene and Quaternary marls, sandstones, clays, and limestones of the Moesian Platform to the north. Along the coast between Varna and Balcik, there has been slumping in soft sediments beneath

Popov V, Mishev K (1974) Geomorphology of Bulgarian Black Sea Coast and Shelf BAC. Publishing House of the Bulgarian Academy of Sciences, Sofia Shuisky YD, Simeonova G (1982) On the types of abrasion cliffs along the Bulgarian Black Sea coast. Eng Geol Hydrogeol 12:11–21 Simeonova G (1976) Coastal processes along the Bulgarian coast. Proceedings of the symposium on dynamics of shoreline erosion, Meziereba, Tbilisi, pp 235–238 Simeonova G (1985) Bulgaria. In: Bird ECF, Schwartz ML (eds) The  World’s Coastline. Stroudsburg, PA, Van Nostrand Reinhold, p 455–458 Vasilev T (1978) Morphodynamics of the Bulgarian coast in stormy waters. Probl Geogr 2:51–55 Zvetkova V, Simeonova G (1984) Accumulation of heavy minerals in landslides in the Burgas area of the Black Sea coast. Mar Geol 54:309–318

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8.26 Romania

Roger Charlier · Constance Charlier

1.  Introduction The Romanian (Rumanian) coast extends for 228 km from the Bulgarian border to the Chilia arm of the Danube River, which enters the Black Sea at the border of Ukraine (Charlier and Piety 1985; Vespremeanu 2004). Tides along this part of the Black Sea coast are semidiurnal, with a maximum amplitude of about 0.12 m. Generally, the tide range is 0.08–0.09 m, and tidal oscillations are masked by meteorological effects, such as variations in barometric pressure and movements by onshore and offshore winds. Water level in the Black Sea is also influenced by exchanges through the Bosporus, precipitation and evaporation, and river discharge, especially from the Danube, which supplies 38% of the river water entering the Black Sea. The effects of river discharge vary seasonally, the lowest runoff being in winter, increasing in spring to a maximum in May, then diminishing to autumn, when there is an increase again due to the rainy season, and declining with winter freezing. Onshore winds, predominantly from the northeast, generate waves from this direction, but the continental shelf is wide along the Romanian coast, and even during storms waves rarely exceed 3 m. The southern part of the coast (>Fig. 8.26.1) from the Bulgarian border to Constanţa has a hilly hinterland and cliffs cut into Sarmatic sandstones and limestones, capped by loess. On headlands these plunge into the sea to depths of up to 60 m. The landslides of the coast of Bulgaria diminish northward. Inland is the Dobrogea Plateau, a dissected ridge from which rivers drain to the coast and to the northward-flowing Danube River on the western side. The lower parts of valleys draining to the coast are occupied by lakes or lagoons impounded by sandy barriers. North of Constanţa the coast is low-lying, with large lagoons (known as limans) enclosed by spits, then the Danube delta (>Fig. 8.26.2). The sea floor adjacent to the Romanian coast is characterised by a broad continental shelf with a uniform slope of 7–12° to a depth of 200 m. The sediment of the Romanian shore are composed of deposits of sand, fluvial mud from the Danube, and deep sedimentary mud mixed with shells. A sandy bottom extends all along the coast in a relatively

narrow zone that extends out from the shore to a depth of 15–25 m. At the mouth of the St. Gheorghe branch of the Danube, this sand zone extends to the 40 m isobath.

2.  The Romanian Coastline Vama Veche, an old customs station, stands just north of the Bulgarian border, and Mangalia is a major seaside resort. To the north is the Romanian Riviera, a chain of modern seaside resorts focussed on beach recreation (>Fig.  8.26.3). Nearshore breakwaters, parallel to the coastline, have been used to shape beaches into a series of bays between artificial tombolos (breakwaters linked to the mainland) (Spataru 1990). Cliffs and bluffs cut into Tertiary and Pleistocene limestone and sandstone rise to 60 m locally along the coast north from the Bulgarian border, notably near Eforie. There are some small limestone capes. Cliffs are receding at up to 7 m/year, and are fronted by narrow beaches. They are interrupted by valley-mouth depressions in which are the remnants of several former lagoons, cut off from the sea by sandy barriers up to 5 m high. There are several lakes, two of which are in hollows completely isolated from the sea and have a limited source of fresh water, partly subterranean. Lake Mangalia, near the Bulgarian border, is surrounded by steep slopes rising 20 m. Thermal mineral water flows from the banks at some points, and is sought for medicinal purposes. Lake Terchirghiol is larger, and occupies a depression extending up to 10 m below the level of the Black Sea. This was formerly a marine gulf, but it became sealed off from the sea by sand deposits, which form the barrier beach of North Eforie. This coastal lagoon became cut off from river and sea inflow, and became hypersaline. The salinity of Lake Techirghiol is extremely high (95.5%), as is the mineral content, and because of the high concentration of minerals and organic material, the black sapropelic mud from this lake is considered therapeutic for numerous ailments. Sheltered from north winds by the surrounding hills, the surface water temperature ranges from 22°C to 26°C in summer.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.26, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Romania

⊡⊡ Fig. 8.26.1 The southern coast of Romania. (Courtesy Roger Charlier and Anita Brandes, prepared for the previous edition of the Encyclopedia of Coasts.) Istria HISTRIA

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There are also coastal lagoons fronted by narrow sandy beaches, which extend as spits and barriers across the mouths of valleys that were submerged by the Late Quaternary marine transgression. In these sectors the beaches are backed by dunes up to 35 m high. The lakes, known locally as ghiols, include Tatlageac, Agigea and Tabacarie. Some have become fresh-water lakes, fed by springs and small streams (Chiriac 1966). The major city of Constanţa (Tomis in Roman times) stands on the narrow Tomis peninsula, and extends north towards Lake Tabacarie and south along the cliff surrounding the harbour. The peninsula, bordered by the Gulf of Tomis, is about 1.5 km long, and has a mean width of 0.65 km. The shore resort of Mamaia, to the north of the city, stands on the barrier that encloses Lake SiutGhiol, which receives abundant fresh water from rivers and springs. The beach is of very fine sand and has been reduced by erosion (Spataru 1990). The sea floor slopes gently seaward: at 100 m from the shoreline, the average depth reaches only 1 m. The width of the beach may be reduced as a result of the development, 5 km to the north of Mamaia, of an industrial petrochemical complex with a long docking jetty which will impede the southward drifting of sand alongshore. Dunes back the shore north of the resort (Mamaia-Sat). The three large lakes north of Constanţa—Lake SiutGhiol, Lake Gargalic and Lake Tasaul, have channels or narrow outlets linking them with the Black Sea. They are shallow, with a depth of 2–3 m, except for Lake Siut-Ghiol, which occupies a former marine inlet at the mouth of the Carasu River. The lake was enclosed by the southward growth of the Mamaia barrier spit. It has an average depth of 4–6 m and trenches reaching a depth of 22 m. Small seasonal variations in lagoon level occur, with the lowest level in summer. The lake is also fed by underground springs. The salinity of these lagoons varies, the lowest being Lake Siut-Ghiol (0.3%) and the highest (1–1.5%) Lake Tasaul and Lake Gargalic. The nutrient balance and oxygen content in each lake are favourable for aquatic flora and faunal, including fisheries. The island of Ovidiu is located in Lake Siut-Ghiol. The northern part of the Romanian coast is low-lying, with extensive sandy beaches and spits, and frequently shifting nearshore bars. The sand has been derived from the fluvial sediment (sand, silt and clay) delivered to the coast by the Danube through its delta distributaries, and carried southward by longshore drifting in response to the predominance of north-easterly waves. North of Cape Midia a former embayment has been enclosed to form a large group of coastal lagoons

Romania

⊡⊡ Fig. 8.26.2 The northern coast of Romania. (Courtesy Roger Charlier and Anita Brandes, prepared for the previous edition of the Encyclopedia of Coasts.)

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8. 26

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A (the Razelm Complex) by a series of spits that have grown southward, the present outer barrier spit being about 60 km long. In the fifth century bc Greeks founded the port city of Histria. It was a coastal settlement on the

shores of this embayment (the former Gulf of Halmyris), but it now stands about 8 km from the sea. The Lakes Sinoe, Zmeica, Galovita and Razelm are shallow (maximum depth 3 m) with a total area of 731 sq. km. They are

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Romania

⊡⊡ Fig. 8.26.3 The recreational beach at Saturn, one of several modern seaside resorts north of Mangalia. (Courtesy Geostudies.)

linked to the Black Sea through four small and variable outlets, the most important of which is Portita (The Little Door). When the lake levels are high, after the inflow of floodwater from the Danube by way of the Dranov and Dunavatu channels, flows out through these openings to the sea, whereas in dry periods there is salt water inflow. The salinity of the lakes varies, with the normal concentration ranging from l% in Lake Razelm, 2–3% in Lake Galovita, and up to 6% in Lake Sinoe. During dry seasons or when storms breach the barrier spit, the influx of seawater augments the salinity, which may then reach 6% in Lake Razelm, 10% in Lake Galovita and 18% in Lake Sinoe. Conversely, inflow of fresh water from the Danube during floods can reduce salinity to 0.2% in Lake Razelm and 1% in Lake Sinoe. There is extensive reed and rush swamp around these lakes, as well as freshwater weed growth, especially in the less disturbed southern parts. The lakes are biologically rich, producing large quantities of fish, and thus contributing significantly to the economy of the area.

3.  Danube Delta The Danube delta occupies an area that was formerly an estuary some 100 km long. Massive sedimentation in Holocene times has built a delta with an area of 4,340 sq. km, of which more than half is composed of ponds, islands, lagoons, and floating rafts of matted roots and

reeds known locally as plauri. Mean annual discharge is nearly 85,000 m3/s. The delta shore is microtidal, but the Black Sea can vary in level by up to 0.40 m as ­meteorological effects accompany tidal oscillations. The delta coast incorporates beach ridges and sand dunes which indicate stages in the evolution of the coastline, and is traversed by natural channels and artificial canals. There are also cheniers, known locally as grindu. Large areas of land are flooded annually, resulting in the gradual raising of the delta plain by alluvial deposition. The Danube delta has distributaries flowing to several river mouths. At each of these the interaction of ­outflowing fluvial currents and sediment with incoming waves and associated (mainly southward) currents in an almost tideless system result in the formation of sand shoals and bars and the dispersal of silt and clay sediment (Giosan et al. 1999). One such bar built up to the point of nearly closing the St. Gheorghe outlet, leaving a depth of only 1.5 m. Successive maps show that the growth of the Danube delta has been accompanied by the waxing and waning of the various distributaries (>Fig. 8.26.4). Dikes have been built to maintain navigable channels. In recent decades most growth has been at the Chila arm, while the Sulina Arm, where breakwaters and levees have been built to maintain a navigable channel, opens to a relatively stable section of delta coast and the St. Gheorghe outlet has been modified by erosion.

Romania

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⊡⊡ Fig. 8.26.4 Growth of the Danube Delta. (Courtesy Geostudies.)

References Charlier RH, Piety JW (1985) Yugoslavia. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, Stroudsburg, PA, pp 439–442 Chiriac V (1966) Contributions Roumaines à 1’étude de 1a Mer Noire. Comité d’Etat des Eaux, Bucarest, Romania Dolotov Y, Kaplin P (2003) Black and Caspian seas ecology and geomorphology. J Coast Res 19(4):194–199

Giosan L, Bokuniewicz H, Fanin N (1999) Longshore sediment transport pattern along the Romanian Danube delta coast. J Coastal Res 15:859–871 Spataru AN (1990) Breakwaters for the protection of Romanian beaches. Coast Eng 14:129–146 Vespremeanu E (2004) Geography of the Black Sea (in Romanian). University of Bucharest, Bucharest, Romania

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8.27 Ukraine

Yurii Shuisky

1.  Introduction From the Rumanian border on the Kilijski branch of the Danube delta, the Ukraine coastline extends some 1,628 km, past the Crimea to Kerchenski Strait, the entrance to the Sea of Azov. Of this coastline, 486 km (29.9%) are erosional, 553 km (34%) stable, and 589 km (39.1%) prograded, but in recent decades some of the prograded sectors have become subject to erosion. The Ukraine coast also extends to the western part of the Sea of Azov, which is covered in a separate chapter (> Sea of Azov). The climate is cool temperate with a cold winter, Odessa having mean monthly temperatures that range from −4°C

in January to 22°C in July and a mean annual rainfall of 406 mm with a summer maximum. The prevailing onshore winds are from SW and SE, respectively, generating eastward and westward longshore drift along beaches, as indicated by the Tendra and Jarylgach spits. Tide range is very small in the northwest Black Sea. In this Ukrainian coastal region, rock formations have been gently folded, and geomorphological evolution has produced alternations of hilly country, valleys, and wider lowlands (>Fig.  8.27.1). Holocene marine submergence initiated development of a coastline with cliffed sectors bordering hilly country, as in the southern part of the Crimea Peninsula (>Fig. 8.27.2), the Tarkhankut Peninsula,

⊡⊡ Fig. 8.27.1 Western Ukraine. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.27, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Ukraine

⊡⊡ Fig. 8.27.2 Crimea. (Courtesy Geostudies.)

and in Odessa and Feodosia bays. Coastal plains and wide valley mouths have been submerged by the Holocene marine transgression to form limans, which are partly or wholly enclosed by shelly sand spits or barriers as lagoons, particularly between the Danube and the Dnieper deltas (Shuisky and Schwartz 1981). Narrower valleys incised into hilly country have been submerged to form rias, as in southwestern Crimea. Faulting has influenced the outlines of the northern Tarkhankut Peninsula (Tkachenko 1970). Where tertiary limestones, sandstones, shales, and volcanic rocks outcrop at the coast, cliffs up to 100 m high have been cut. These have been retreating at rates of up to 30 cm/year during the past century. Softer formations, such as clays and loams, have formed steep sectors subject to recurrent landsliding (>Fig. 8.27.3), and on these recession has been up to 9 m/year. On depositional coastlines, beaches, spits, and beach-ridge plains can gain or lose up to 20 m in a single storm, but over longer periods the rates of change are similar to those on soft cliffs (Shuisky 1974; Shuisky and Schwartz 1980). The coast north of the Danube delta is smooth and low-lying, formed by the deposition of sand, silt, and clay

derived from Danube river sediment. It has prograded up to 10 m/year. The Dniester and Dnieper have built deltaic plains into former limans, rather than protruding deltas. Where sandy sediment have been deposited, beaches are backed by derived dunes generally 2–2.5 m high, occasionally reaching up to 6 m, in a backshore ridge 30–40 m wide extending intermittently along the coast. Prevailing northerly winds drift sand southward, which on many sectors is seaward, so that large-scale dune development is impeded. Shell sands, mainly mussels, are an important constituent of marine sediment, typically up to 5%, but often more; some beaches consist entirely of molluscan material. Beaches and dunes are dominated by fine sandy material, silt and clay being dispersed seaward. The supply of sandy sediment from rivers, from cliff erosion, and from the sea floor has diminished in recent centuries, and at present only 7.2% of the Ukrainian coastline is actively prograding. Longshore drifting has produced the major barrier spits of Tendra and Jarylgach, the cuspate spit Bakal, and many coastal barriers. Between Odessa Bay and Zhebrianska Bay, the longshore sediment flow is over 100,000 m3/year, between the Dnieper liman and Odessa

Ukraine

⊡⊡ Fig. 8.27.3 Large slump block on a cliff near Port Yuhzny, east of Odessa. (Courtesy M.L. Schwartz.)

Gulf about 20,000 m3/year, and between Khersones Cape and Kalamitski Gulf over 70,000 m3/year.

2.  The Ukraine Coastline The northern lobe of the Danube delta protrudes into the Ukraine. To the north is a long smooth coast of spits and barriers extends to Odessa, enclosing a series of valleymouth limans, including Oz Sasyk, Oz Shagany, Oz Alibey, and the Dnezstrovskiy Liman. There is also an elongated narrow lagoon south-west of Bugaz, where a segment of the low-lying coastal plain has been submerged and bordered by a sandy barrier. Odessa has a large artificial beach. The Yuzhnyy Bug flows south into Bugskiy Liman, and the Dneiper River flows westward into the head of Dneprovskiy Liman. On the

8.27

Pokrovskiy Peninsula, there are many lakes. To the south, the long barrier island, Tendrovskaya Kosa, is backed by the Tendrovskiy Zaliv lagoon. On the Tarkhankut Peninsula (western Crimea), steep eroding coastal slopes in loess are retreating at up to 20 m/year (Shuisky 1974). Steep receding and slumping cliffs border the mountainous coasts of the southern Crimea. The town of Alupka is built on landslide topography, the stability of which is influenced by groundwater seepage, basal cliff erosion, and loading by buildings. Coastal slopes in the Karadag region are on old volcanic deposits. Volcanic rocks and limestones form headlands between bays cut out in shales, clays, and sandstones. There are rias near Sevastopol and Balaklava. Beaches on the south coast of the Crimea are mainly of shingle, the finer sediment having been lost offshore. Extraction of gravel from beaches has resulted in erosion, but recently some beaches have been artificially restored. Coastal erosion has resulted in extensive protection by sea walls, breakwaters, and groynes with intervening artificial beaches.

References Shuisky YD (1974) Processes and rate of abrasion of the Ukrainian shoreline of the Azov and Black Seas. P Geogr Soc USSR 6:117–125 Shuisky YD, Schwartz ML (1980) Processes of development of eroding  and slumping shores on the Black Sea coast. Shore Beach 48:36–39 Shuisky YD, Schwartz ML (1981) Dynamics and morphology of barrier beaches of the Black Sea coast limans. Shore Beach 49:45–50 Tkachenko GG (1970) Influence of tectonic movements on coastal morphology. In: Tkachenko GG et al. (ed) Geology of the coastline and bottom of the Black and Azov Seas, vol. 4. pp 27–33

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8.28 Sea of Azov

1.  Introduction The western part of the Sea of Azov is in Ukraine and the eastern part in Russia. This account will deal with the Sea of Azov as a whole, then describe and illustrate the Ukrainian and Russian coastlines. The Sea of Azov has an area of 38,000 sq. km and is shallow, with a maximum depth of only 15 m. It occupies a basin that originally formed in Miocene times, and was periodically submerged by the sea during Plio-Pleistocene times, attaining its present form as the result of Late Quaternary marine submergence of an undulating lowland. It is bordered by hilly country and tributary valleys of inflowing rivers, notably the Kuban and the Don, as well as the Protoka, Yeya, Mius, Molochnya, and some smaller streams. On the southern side, the Kerch Peninsula on the eastern part of the Crimea and the Taman Peninsula are separated by Kerch Strait, a narrow connection with the Black Sea. There are many spits and barriers of shelly sand, backed by lagoons and marshes (>Fig. 8.28.1). The climate is temperate continental, with cold winters  when E and NE winds predominate and extensive nearshore areas freeze. In the summer months, there are westerly and southeasterly winds. Wave action therefore depends on coastal aspect in relation to these wind re­­ gimes. Salinity is generally low, but varies with the interaction of brackish water entering through Kerch Strait and the varying quantities of rain and river input. The Sivash Lagoons, behind the Arabat barrier in the southwest, are generally brackish. Tide ranges are small, but sea level fluctuates by up to 0.33 m as the result of variations in rain and river inflow and wind stress, periods of strong wind action being followed by oscillatory seiches. Currents are generally counterclockwise, but occasionally reverse. There is high biological productivity in the shallow, generally well-mixed sea water, which is warm in summer and receives abundant nutrients from the inflowing rivers. Seagrasses, notably Zostera, are dense in shallow water, and fish and shelly organisms are abundant.

The coastline of the Azov Sea measures 1,860 km, of which 420 (22.6%) are erosional, and 1,440 (77.4%) have prograded. Within the latter, only 141 km are actively prograding beaches and 46 km advancing deltas; the remaining 1,299 km were formerly prograded and are now undergoing erosion. In the Russian sector, the Don River discharges nearly 1.7 million m3 of sediment annually (to the Taganrog Gulf), and the Kuban River 2.6 million m3 (to the Yasen and Temryuk Gulfs), the remaining streams being regulated with almost no sediment yield (Mamykina 1978). Of a total fluvial sediment yield of 4.5 million m3, 70% is retained in lagoons and on delta shores. The Sea of Azov generates about 25 million tonnes of molluscan shells per year, of which more than 15% (about 1.84 million m3) is added to coastal beaches, spits, and barriers. Most depositional coasts are extremely low-lying (up to 2.0 m) and formed mainly (up to 90%) of shell detritus washed in from the sea floor. Changes in water salinity resulting from variations in river discharge have influenced the productivity of molluscs. The dominant beach-forming mollusc, Cerastoderma glaucum, abounds when the salinity is 11–13 ppt, but during periods of higher or lower salinity this species is replaced by other molluscs with much lower productivity (notably Mytilus galloprovincialis and Dreissena polymorpha). Salinity in the Azov Sea has varied during recent decades from 10 to 14.5 ppt, mainly due to changes in the amount of water extracted from rivers (Kuksa 1994): when it exceeded 13  parts per thousand Cerastoderma declined, and the supply of shelly material to beaches diminished, but since the early 1990s salinity has generally decreased because of  reduced water extraction in the basins of the Don and  Kuban rivers, and there has been revival of shelly deposition. Cliffs are mainly cut in Tertiary and Quaternary clays  and loess, and are receding at rates of up to 6 m/ year.  Limestones retreat more slowly, up to 0.4 m/year (Shuisky 1985). Sediment from eroding cliffs amounts to

Edited version of a chapter by Yurii Shuisky in The World’s Coastline (1985: 467–472); A. O. Selivanov (Lomonosov Moscow State University) compiled the Russian section. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.28, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Sea of Azov

⊡⊡ Fig. 8.28.1 The coastline of the Azov Sea: 1 – receding cliffs, 2 – abandoned cliffs (bluffs), 3 – receding coastlines that formerly prograded, 4 – stable coastlines, 5 – advancing (prograding) coastlines, 6 – direction of longshore drift, 7 – delivery of sediment from the sea floor, 8 – delivery of sediment from rivers and cliff erosion. (Courtesy Geostudies.)

about  6.9  million m3/year, of which 0.85 million m3 are sand coarser than 0.1 mm grain size. Thus, much of the sediment is fine-grained, and quickly lost seaward to the coastal slope, so that it does not form extensive coastal depositional landforms. A further 4.6 million m3/year of sediment is generated by sea-floor abrasion, and of this 0.88 million m3 is sand and gravel. Thus, only a small proportion of the material produced by cliff and sea-floor erosion is suitable for beach accretion. Cliffs have formed where the sea intersects uplands, and elongated bays occupy drowned lowlands and valley mouths. The larger rivers have built deltas, and there has been extensive deposition of both inorganic sediment from rivers and eroded cliffs, and biogenic material, mainly molluscan sand and gravel, washed in from the sea floor to form beaches, spits, and barriers that enclose lagoons.

2.  The Ukraine Coast of the Sea of Azov The western shore of Kerch Strait and the northern coast of the Kerchenskiy Peninsula are embayed, with hilly promontories. To the west is the long gently curved Arabat barrier, a deposit of sand and shells shaped by waves from the ENE, which are slightly refracted as they approach the shore. Behind the Arabat barrier are the Sivash Lagoons, with an irregular inner coastline of rias between narrow peninsulas. The lagoons are separated from Karkinitskiy Bay to the west only by the narrow isthmus of Perekop. The northwest coast of the Sea of Azov is notable for the five large sand spits that have been described and discussed by Zenkovich (1967). They are up to 40 km long and 1–2 m high, slightly recurved, narrowing from 3–4 km wide at the coast. They consist of sediment that has come from the River Don and from the erosion of cliffs,

Sea of Azov

and drifted westward along the coast. They have grown out southwest, at an angle to the coast, as the result of onshore and longshore drifting, shaped in relation to waves arriving obliquely from the southeast. Their southeastern shores are steep and receding as they migrate westward. Secondary accumulation of shelly sand on the leeward (western) side forms a series of curved beach ridges that are truncated as the eastern flank is cut back by wave action. There are broad intervening bays, backed by cliffs and receiving small rivers. To the east is the Russian border.

3.  The Russian Coast of the Sea of Azov The Russian coast of the Sea of Azov has a number of gulfs and bays between large promontories. The NE coast has cliffs up to 15 m high, cut in Neogene Scythian clays and Sarmatian limestones overlain by Quaternary fluvial and lacustrine sand and gravel, and a generally thick layer of subaerial loess intercalated with ancient soils. The cliffs are receding at a rate of 0.5–0.8 m/year, rarely up to 1.5 m/ year. During the Holocene marine transgression, cliffs retreated very rapidly, and remnants of them are found on the sea floor as far as 8 km offshore. The sea floor slope formed during this transgression is of gentle gradient. Storm surges have played an important role in coastal evolution. Most coastal barriers, spits, and deltaic areas are partially or wholly inundated during storm surges of up to 2 m. Cliffs cut in soft Pleistocene loess and marine sediment and Pliocene clays retreat dozens of metres during these events. Storm surges of over 3 m occurred in 1914 and 1969, and resulted in severe coastal erosion. Immediately west of Taganrog, limestones outcrop above the present sea level and there is a shore platform, in places with narrow boulder and gravel beach. In the northeastern corner of the Taganrog Gulf, coastal slopes rise to 50 m in elevation and have extensive landslides. The Taganrog Gulf extends ENE to the delta at the mouth of the River Don. The delta is extensive, with a coastline nearly 40 km long and an area of about 540 sq. km. Coastal marshes occupy bays between the numerous deltaic spurs with channels through which discharge is highly variable. Elevations in the delta do not exceed 2 m. Most of the delta is submerged during spring and summer floods. According to archaeological and historical data, the Don River delta coast has advanced seaward by 8 m/year during the past 2,500 years. At present, the river discharge is distributed between three main channels and many sub-channels and as a result of intensive agricultural

8.28

­ evelopment and water engineering projects in the river d basin, the delta coastline is no longer advancing. Storm surges that occur once or twice a century raise water level up to 3 m and result in widespread inundation, including parts of the large Azov City. Sediment eroded from cliffs and sediment supplied by coast of the Don River and a number of smaller rivers on the north coast has drifted alongshore to form the spits described by Zenkovich (1958, 1967). Most of them are situated in Ukraine but there are several in the Taganrog Gulf in Russia, including Beglitskaya, Petrushina, and Kurichya, generally composed of sands, gravel, and pebble with relatively low carbonate content (less than 30%). Radiocarbon dating has shown that these spits are between 2,000 and 3,000 years old, and as there is archaeological and historical evidence that they did not exist by 500 bc they probably began to form about 2,300 years ago. The spits consist of Holocene sediment up to 2 m thick, their crests up to 2.5 m above sea level. The Holocene sediment are underlain by Neogene clays of similar to those outcropping in cliffs in the intervening bays. The southern coast of the Taganrog Gulf has high eroding cliffs in loess-like Pleistocene sediment. Their elevation reaches 40–50 m at near Stefanidinodar and Margaritovo, and falls to 20 m near Eisk and 3–4 m at the beginning of the Dolgaya Spit. Cliffs are retreating at a rate of up to 4 m/year in embayments and up to 8 m/year on salients, as near Eisk (Mamykina and Khrustalev 1980). Cliff recession is accompanied by landslides, particularly active during storm surges. Where Scythian clays outcrop, there are wave-cut notches near present sea level. There are narrow beaches of fine sand beneath the cliffs, and several small sand spits, mainly extending from headlands, as at Ochakova, Glafirova, Sazalnik, and Eisk. Most of the spits have re­­ceived sediment partly from the Don River and partly from cliff erosion. The content of shelly material increases westward from 30–40% to 80%. Sediment supply from the Don River and from cliffs is now insufficient to maintain beaches and spits, which are diminishing, apart from the Chumbur Spit, formed in a small liman (>Fig. 8.28.2). Eisk Gulf, the largest embayment in the area, is bordered by relatively stable coastal slopes in Pleistocene sediment. In the middle of the twentieth century, 5 million tonnes of sediment were extracted from this coast, partly from the eroded Eisk Spit that separates the gulf from the sea. Only a short segment of this spit, 5 km long and 1.5 m high, and a small island, survive, and the inner part of the spit, repeatedly inundated by storm surges, is defended by groynes and revetments.

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Sea of Azov

⊡⊡ Fig. 8.28.2 The Chumbur Liman on the southern coast of the Taganrog Gulf. (Courtesy A.V. Porotov.)

Dolgaya Spit is a triangular foreland at the south-­ western end of the Taganrog Gulf. It is now over 10 km long and 4 km wide at its inner end. The spit has formed by the accumulation of shells washed in from the sea floor and, to a lesser extent, by alongshore sediment supply from the Taganrog Gulf and the coast to the south. The spit rises 2.5–3.5 m above sea level and is primarily composed of shells and shelly debris increasing from 50–60% at the landward end to 95–98% to seaward. Excavation of this material in the 1960s and 1970s resulted in the drastic retreat of the coastline, particularly toward the end of the spit, which shortened by over 4 km from 1950 to 1975. Subsequently, the prohibition of extraction and increased shell production with diminished sea salinity has led to partial restoration of this spit, but it has not yet reached its natural dimensions. The 35 km coastal segment between the Dolgaya and Kamyshevatian spits is a relatively straight coastline cut in Pleistocene loess-like sediment forming cliffs up to 15 m high, fronted by narrow (15–20 m) beaches. The cliffs are retreating 2–3 m each year, and the reduction of beaches is  leading to accelerated erosion. Kamyshevatian Spit, nearly 6 km long and over 3 km wide at its landward end, is ­composed almost entirely (over 95%) of shelly material, and is sensitive to changes in shell productivity and beach quarrying. At present, the seaward shore of the spit is retreating at a rate of over 5 m/year. South of the Eisk Peninsula, Yasen Spit separates Yasen Gulf from the Beisug Liman. The spit is 18 km long from north to south and just over 1 m high, composed of shelly debris. It is retreating by over 3 m/year and is really a

s­ emi-artificial barrier for Beisug Liman, a brackish water ­reservoir used for fish breeding. The coast of Beisug Liman and the Khan Lake to the north are relatively low slopes in loess-like sediment with alluvial-marine plains in bays. Relict coastal barriers are composed of shell detritus and are relatively stable. There are numerous lakes and marshes in the hinterland. South from the Yasen Gulf for 40 km are cliffs up to 11 m high cut in Scythian Clay, with extensive rock falls and landslides. These cliffs are retreating at up to 3 m/year, except where they are defended by revetments and groynes at Primorsko-Akhtarsk. The Kuban River enters Temryuk Bay in the southeastern corner of the Sea of Azov. This was a large marine gulf until the beginning of the Late Holocene (2,700– 2,300  years bp), but subsequently the Kuban River has flowed through a series of channels and built delta segments. At first, much of river discharge flowed to the Kiziltash Liman and other limans on the coast of the Black Sea, forming a series of small deltas. Later, in the Classical Greek period (2,500–2,000 years bp), the present Taman Gulf served as a main outlet for the Kuban River. At the end of the 1st millennium bc, owing to tectonic deformation and infilling of previous river channels in the southern part of the delta, the greater part of river discharge moved to northern deltaic channels. Until the mid-eighteenth century, most of the river outflow entered the Sea of Azov through the Akhtanizov and Kurchan limans in the SW part of the present delta near Temryuk and through the Protoka River in the north-east. In the late nineteenth century, extensive engineering works on river flow redistribution between the channels

Sea of Azov

was carried out by Cossacks, and resulted in the present Kuban River outflow. As a result of these changes, there is an alternation of the Yasen/Beisug Gulf, Achuevo Gulf, and Temryuk Gulfs with low-lying deltaic peninsulas. In the twentieth century, there was a drastic decrease in sediment supply from the Kuban River due to irrigation works and the construction of reservoirs upstream. It has been estimated that sediment yield fell from 7.9 million tonnes/year in the 1930s–1960s to 0.84 million tonnes/year in the 1970s. The area of the Kuban River delta is 4,300 sq. km, with a coastline over 150 km consisting of barrier spits, usually 2–3 m, rarely up to 4–4.5 m high, separating brackish limans from the sea. These coastal features consist primarily (50–80%) of shelly sand, the proportion of terrigenous sediment increasing to 50–60% at the mouths of active deltaic channels. The shelly sand has been swept in from shoals at 5–7 m depth offshore, especially during storm surges (Zenkovich 1958), but this source has diminished during the past 40 years. A significant proportion of river sediment is deposited in limans, which become mudflats after several decades. Since the eighteenth century, when economic development of coastal area began, natural channels were deepened and artificial canals (girlo) were cut to avoid this infilling of the limans, conducting sediment to form new deltaic protrusions at the coast. Behind the coastal barrier spits and limans is swampy land (plavni), with remnants of shelly sand beach ridges 2.5–3 m in elevation, marking stages in progradation during the Holocene. In places, these beach ridges are now 45–50 km inland. During storm surges and river flooding, much of this deltaic lowland is inundated. During such extreme events in 1914 and 1969, the water level rose more than 3.2 m above mean level and an area 15–18 km wide was inundated. Now nearly 60% of the outflow enters the sea through the Kuban River and its several channels, the rest going through the Protoka River (30%) and the artificial Akhtanivov and Kiziltash liman canals. Changes in river outflow, combined with tectonic ­subsidence, have resulted in extensive shoreline erosion on several parts of the delta coastline and some local ­progradation. Achuevo Spit, north of the Protoka channel, is retreating 5–7 m/year in a sector that was prograding at a rate of 5–6 m/year before a decrease in river flow in the Protoka channel. In the early twentieth century, the Petrushin Channel of the Kuban River, still the primary one nowadays, received over 70% of the total river sediment and the delta coastline advanced by 35 m/year, but a drastic decrease of river sediment discharge due to irrigation and, especially the construction of reservoirs in the middle reaches of the Kuban River, resulted in the lower rates of coastline advance, 3–5 m/year, during the

8.28

past decade. The seaward edge of the barrier separating the Akhtanizov and Kurchan limans is retreating 1.5–2 m/year. In general, the delta area has diminished during the past few decades, largely as a result of human activities. The SE coast of the Sea of Azov consists of an alternation of capes in hard Neogene limestone, as at Pekly, Kamenny, and Akhilleon, and embayments cut in Neogene and Quaternary sediment. The capes have a complex geology, with calcareous and marly sandstones and a cover of loess-like and clayey Quaternary sediment. Landslides have occurred where clays alternate with sandstones, and cliffs have retreated by up to 9 m/year. Mud volcanoes abound on the surface near the coastline, adding to its destruction. Pre-Quaternary rocks outcrop locally, forming residual rock formations with tombolos (>Fig. 8.28.3). There are narrow sand beaches in bays backed by slopes in Quaternary sediment. Kerch Strait connecting the Sea of Azov to the Black Sea was the route taken by the Don River during Pleistocene low sea level phases, submerged during the Holocene marine transgression. The mean depth of the strait is 5 m and its width varies from 9 km at Chushka Spit to 25 km at the Taman Peninsula. The east coast of Kerch Strait is generally low-lying. The southward-oriented Chushka Spit, nearly 17 km in length, diverges from Akhilleon Cape, the westernmost point of the Russian Sea of Azov coast. The spit has been fed from coastal and sea floor erosion in this area, but now is receiving very little sediment, and a number of gaps have been cut through it during storm surges. The spit is defended only by primitive structures such as concrete blocks and the bulkheads that partially support the Caucasus to Crimea railway, which runs along the spit for 17 km. The former Tuzla Spit, earlier fed from both the Sea of Azov and the Black Sea, has been dissected and reduced to a small residual island in Kerch Strait less than 1 m in elevation, with a shoal to the east where the spit formerly extended. The coasts of the Taman Gulf on the eastern side of Kerch Strait are dominated by lowlands that are submerged during storm surges. Their flat surfaces composed of fine sand and silt, and there are various kinds of mud volcano in the area. To the south, the Taman Peninsula is of Neogene limestone, up to 160 m in elevation and bordered by extensive loessic and clayey Quaternary sediment exposed in high cliffs. The geological sections are the most important source of information on Quaternary history in this area. The peninsula has the sites of many Classical Greek towns, usually on relatively low unstable coasts. Cliffs cut in Neogene limestones retreat several cm/year, whereas those in Pleistocene loess-like sediment are receding up to 1 m/year.

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Sea of Azov

⊡⊡ Fig. 8.28.3 An embayed cliffed coastline on the southern coast of the Sea of Azov east of Kerch Strait. (Courtesy S.L. Nikiforov.)

References Kuksa VI (1994) Southern Seas under the anthropogenic impact. (in Russian). Gidrometeoizdat, St. Petersburg Mamykina VA (1978) Recent processes in the coastal zone of the Azov Sea. P Geogr Soc USSR 110:351–359

Selivanov AO (2001) Coastal disaster on the Sea of Azov: myth or real threat? (in Russian). GEOS, Moscow Zenkovich VP (1958) Coasts of the Black Sea and the Sea of Azov. (in Russian). Geografizdat, Moscow Zenkovich VP (1967) Processes of coastal development. Oliver and Boyd, Edinburgh

8.29 Republic of Georgia

Pavel Kaplin · Andrei Selivanov

1.  Introduction The coastline of Georgia extends southeast from the Russian border near Adler to the Turkish border south of Batumi, a distance of 312 km. The coastal features are determined by several factors: the coast-parallel structure of the Great Caucasus alternating southward with the large synclinal Colchis Lowland and the Little Caucasus range near Batumi, the narrow and steep sea floor, which allows large waves to reach the coast, the generally southward longshore drifting, and the massive loss of sediment into  submarine canyons that cross the shelf in front of the  mouths of the rivers Mzymta, Psou, Kodori, and Chorokhi. In the latter half of the twentieth century, the fluvial sediment supply to the coast was diminished by construction of dams in the larger rivers. There was also large-scale extraction of sand and gravel from beaches, totalling 30 million m3 in the 1940s–1970s (Kiknadze 1977). The natural coastal system was further modified by the construction of solid coast defence structures, including sea walls, revetments, groynes, bulkheads, and breakwaters, as well as port jetties as at Poti and Batumi. As a result, 155 km (49.6%) of the Georgian coastline retreated by up to 3 m/year in the early 1960s, and by the early 1980s, this had increased to 220 km (70.5%) of the coastline. Between the early 1980s and the early 1990s, artificial beach nourishment, using over 12 million m3 of sand and gravel, restored many beaches, particularly in resort areas (Zenkovich and Schwartz 1987). In the northern part of the Georgian coast, sediment discharge from the Mzymta and Psou rivers has built a wide coastal promontory since the Middle Holocene. Eventually, the prograding coastline approached the heads of submarine canyons, and much sediment was lost into them, but part of the longshore drift (20,000–30,000 m3/ year) proceeds further south. Further to the southeast are the extremely steep coastal slopes of the Gagra range, an offset of the Great Caucasus composed of resistant Cretaceous and Jurassic limestones. The sea floor is fairly steep and waves are reflected from resistant underwater outcrops, so that sediment generally

move offshore and are deposited at a depth of 3–6 m. There are generally no beaches, except locally, retained between groynes in embayments. In several river valleys, there are depositional terraces 3–10 m above present sea level, with Pleistocene marine molluscs indicating that the sea then intruded into the valleys. Gagra had a wide beach until erosion started after the construction of a 150 m jetty to the west of the resort in 1914. The beach then prograded west of the jetty until longshore drifting to the Gagra beaches resumed. In general, there has been beach erosion as the result of the reduction of longshore drifting from the Russian Black Sea coast to the northwest, and because of large-scale extraction of sand and gravel for construction purposes. Groynes were built in an attempt to retain beach sediment, but erosion is dominant along the whole coastal segment south to Bzyb River. Only where beach nourishment has taken place, as in Gagra Bay, have beaches been partially restored. Southeast of the mouth of Bzyb River is the Pitsunda Peninsula with its unique pine-grove landscapes and tall hotels. During the 1960s, the coast of the peninsula progra­ ded to the heads of submarine canyons. Some 200,000 m3 of sediment from Bzyb River drifted SE to the head of the ­peninsula annually. Most are lost into submarine canyons, the Akula (Shark) canyon receiving over 80,000 m3 of ­sediment per year (Kaplin et al. 1993). In the bay south of the cape, the submarine coastal slope is very steep, descending from depths of 10–12 m at 25–30° to a flat shelf surfaceat 70–90 m. The coast here is eroding, with and cliffs and shore platforms cut in Cretaceous limestone. There are only narrow gravel beaches, supplied by local northward sediment movement. Solid coastal defence structures were built, together with the hotels, in the late 1960s and the early 1970s, and resulted in destruction of the coast. In the late 1970s, several groynes and jetties were demolished and extensive  beach nourishment began, forming wide beaches (>Fig.  8.29.1). SE of the Pitsunda Peninsula, steep cliffs are cut in soft stratified sediment (>Fig.  8.29.2). At the mouth of the Belaya River, fluvial sediment moves southward across a submerged delta at about 15,000–20,000 m3/ year. This is augmented on a smaller scale by fluvial

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_8.29, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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8. 29

Republic of Georgia

⊡⊡ Fig. 8.29.1 Cape Pistunda, showing the artificial beach in the early 1990s. (Courtesy S.L. Nikiforov.)

s­ ediment drifting southward from the Gumista River. A cape formed both by the river and longshore drift from the north is similar to the Pitsunda Peninsula, and diminishes significantly the sediment supply to Sukhumi Bay. Sukhumi Bay is backed by a wide coastal lowland with a series of Quaternary river terraces. Wide beaches (up to 100–120 m) were much reduced after the construction of groynes, but partially restored by later artificial beach nourishment. Further south lies the Kodori River delta, which also forms an obstacle for longshore drifting, while much sediment is lost into the submarine canyons at Iskuriya Cape. For at least 10 km south of Kodori Cape, the Holocene deltaic terrace is now eroding rapidly. The principal embayment in the eastern part of the Black Sea, nearly 90 km in length, adjoins the extensive Colchis (Kolkhida) Lowland. It is an area of tectonic subsidence proceeding at the rate of more than 6 mm/year in the central part of the lowland. In the middle Holocene (5,000–4,000 years bp), the Colchis Lowland was a large gulf extending 70–90 km inland from the present coast. High sediment discharge from rivers in the Great and Little Caucasian ranges (the largest being the Rioni River) filled this gulf in several millennia and formed the present lowland lying only a few metres above sea level. Pleistocene deposits reach 500 m in thickness and are composed generally of intercalating lagoon peats and beach sands, the latter elongated parallel to the present coastline and possibly representing relict Pleistocene coastal barriers. The present coastal barrier separates a vast swampy area from the sea. Sandy beaches along the Colchis coast

are not wide, and in most places lagoon and lacustrine clays exist behind them. Construction of dams and various types of river flow redistribution in the Rioni and Inguri Rivers have modified the coastline, resulting in rapid retreat in some places and rapid advance in others. In the northern portion of Colchis Lowland, the coast is generally retreating at over 3 m/year, but in the south it is relatively stable. Submarine canyons lie close to the mouths of large rivers, such as the Inguri, Rioni, and Supsa, and much of the fluvial sediment yield is lost into these canyons (Zenkovich 1985). The Rioni River formed a smooth cuspate delta, which prograded until the 1920s when it reached the head of the submarine canyon. Since that time, the canyon has deepened and its head has moved toward the shore. In 1939, the river migrated to a new channel 6 km north of the old one, and thereafter the delta grew rapidly  at the new site, while the former delta was rapidly eroded, the coastline retreating by about 900 m by the middle 1970s. Simultaneously, the canyon head ceased to  move shoreward and is now gradually filling with sediment. Coastal sectors near Kobuleti and Batumi in the southern part of the Colchis Lowland have wide gravel beaches, which survive notwithstanding the general decrease of sediment supply from rivers. However, slow erosion of beaches is evident. South of the Colchis Lowland is the Chorokhi River delta, which is similar to the other deltas from which sediment is lost into submarine canyons. There are two of them in this area.

Republic of Georgia

8. 29

⊡⊡ Fig. 8.29.2 Eroding cliffs SE of Cape Pistunda. (Courtesy Geostudies.)

Erosion is prevalent along the coast south of the delta to the Turkish border.

References Kaplin PA, Porotov AV, Selivanov AO, Yesin, NV (1993) The North Black Sea and the Sea of Azov under a possible greenhouse-induced

s­ ea-level rise. In: Kos’yan R, Magoon OT (eds) Coastlines of the Black Sea. American Society of Civil Engineers, New York, pp 316–354 Kiknadze AG (1977) Dynamic schemes and sediment budget of the Georgian coast on the Black Sea. In: Davitaya FF (ed) Man and environment. (in Russian). Alashara, Sukhumi, pp 5–11 Zenkovich VP (1985) The Eastern Black Sea, USSR. In: Bird ECF, Schwartz M (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 467–472 Zenkovich VP, Schwartz ML (1987) Protecting the Black Sea - Georgian S.S.R. gravel coast. J Coastal Res 3:201–209

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9.0 Russian Federation – Editorial Introduction

The Russian Federation has long Arctic and Pacific coastlines and shorter sectors on the Baltic, Black, and Caspian Seas. The Arctic and Pacific coasts are influenced by the broad geological pattern, while the short sections of Baltic and Black Sea coast and the N W Caspian Coast are where the Great North European Lowland is intersected (Zenkovich 1967).

1.  Baltic, Black Sea, and Caspian Coasts The Russian East Baltic Coast borders the easternmost part of the Gulf of Finland, with St Petersburg at its head. To the south is > Kaliningrad, between > Lithuania and > Poland. These are low wave energy, almost tideless coasts beside a fresh water sea that freezes in winter. The > Russian Black Sea Coast includes part of the Sea of Azov, the Crimean Peninsula, and the western Caucasus. The Black Sea is almost tideless, but storms can generate strong wave action and occasional storm surges from the SW. The coastal outlines are much influenced by the pattern of geological outcrops. In the southern part of the Crimea Peninsula, the mountains have resistant Jurassic limestones over weaker Triassic shales and on coastal slopes there are landslides where the limestone caprock has collapsed and disintegrated to scattered rocks on slopes of Trias. The Russian sector of the Caspian coast (> Russian Caspian Coast) borders a broad lowland, including the large Volga delta, extending south to the flanks of the Cau­ casus Mountains. The climate is semi-arid, and the coast has been subject to major fluctuations of sea level, with a marine transgression that started in 1977. The tide range is small, but wave action can be strong from the SE.

2.  Pacific Coast The > Pacific Coast of Russia has varying degrees of exposure to ocean waves because of the sheltering effects of Hokkaido, Sakhalin, and Kamchatka. The climate varies

from temperate in the south to subarctic and arctic in the north, with a consequent increase in shore and sea ice. Much of the coast is steep and cliffed, and the beaches are often gravelly, fed by frost-shattered cliffs and short, steep rivers. Steep coasts and bedrock headlands on metamorphic and igneous rocks occur on the Chukchi Peninsula, and on either side of the Anadyr Gulf, which has shingle beaches and barriers. There is extensive shore ice in winter. To the south, the steep coast continues through eastern Kamchatka, with four large volcanic promontories (up to 800 m high) and narrow lowlands with extensive shingle beach ridges behind bays. This coast is subject to Pacific tsunamis as well as local earthquakes. The Kuril Islands, a chain of islands of lava and volcanic ash, extend south from Kamchatka toward Hokkaido in Japan. Western Kamchatka has a broader coastal plain crossed by several rivers, and has extensive sand and shingle beaches and barriers, some of which deflect river mouths. The tide range is generally small ( Arctic Coast of Russia) has features related to Pleistocene glaciation (with ice persis-ting locally, as in Novaya Zemlya), postglacial isostatic uplift, deep permafrost, extensive sea and shore ice in the long winter, wave action in the brief summer, and ­generally small tide ranges (except in the west, where it locally exceeds 3 m). The coast is irregular, with periglacial

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Russian Gulf of Finland

features such as tundra cliffs interspersed with rocky promontories. There are peat cliffs and tidal marshlands. It is icebound for much of the year. Russia includes the eastern part of the Great North European Lowland, west of the Ural Mountains. In the northern part of this lowland, the crystalline Fennoscandian Shield borders the White Sea and the Barents Sea, forming rocky coasts and sectors of steep, plunging cliff interrupted by fiords. To the west, the Kola Peninsula has headlands of granite gneiss between bays with lowlands occupied by glacial drift deposits. Glacial drift is also extensive south and east of the White Sea, where permafrost produces tundra cliffs that show slumping and erosion during the summer thaw. The mean spring tide range in the Mezen Gulf exceeds 10 m. To the east of the White Sea, the Kanin Peninsula consists of Palaeozoic rocks projecting into the Barents Sea, with low cliffs in glacial drift to the east. The Urals are a mountain range of strongly folded PreCambrian and Palaeozoic formations that run N-S from Novaya Zemlya in the north to decline beneath sedimentary formations toward the Aral Sea. On Novaya Zemlya, there are residual glaciers, and the steep coasts are trenched by many fiords and fringed by strandflats. East of the Urals is the broad West Siberian Plain, a lowland drained by the Ob and Yenisey Rivers that flow north into long inlets that open to the Kara Sea. The coastal region is low-lying, with deep permafrost, and cliffs of peat or glacial drift that retreat rapidly during the ­summer thaw. To the east, the Taymyr Peninsula has glaciated uplands,

the Burranga Mountains, on Pre-Cambrian ­metamorphic rocks that extend locally to the coast. These are a northeastern branch of the Urals that continues on to the Severnaya Zemlya islands. On their southern side, the Kjeta River drains the Taymyr Depression, opening to a long estuary and the Laptev Sea. East of the Yenisey River is the Central Siberian Platform, which declines northward to the North Siberian Plain, another permafrost area with tundra bluffs exposed to waves from the Laptev Sea during the brief summer thaw. The Lena flows down to a major delta. East of the Lena mountain ranges extend from the Buorkhaya Bay southward and eastward to the Sea of Okhotsk. They are bordered northward by the swampy coastal plain of NE Siberia, bordering the shallow Laptev and East Siberian Seas. Again, this is a permafrost region, with tundra bluffs cut in glacial drift and long shingle beaches, spits, and barriers. East of the Kolyma River, the coastal plain narrows, and spurs of the hilly Anadyr hinterland end in bedrock headlands between bays with rapidly receding tundra bluffs. Beyond Chaunskaya Bay, the hinterland is more mountainous and there are steep rocky coasts alternating with narrow swampy lowlands bordering the Chukchi Sea and extending to Dezhneva Cape, beside Bering Strait.

Reference Zenkovich VP (1967) Processes of coastal development. Oliver and Boyd, Edinburgh

9.1  Russian Gulf of Finland   > 8.3.1

Russian Gulf of Finland

9.2  Russian Black Sea Coast   Pavel Kaplin · Andrei Selivanov  

Introduction

small dunes. The barrier is composed of ancient alluvial sediment from the Kuban River, which until the late 19th The north coast of the Black Sea within the borders of century discharged into the north Black Sea limans. Now Russia extends for nearly 300 km southeast of Kerch Strait, the northern part of the barrier, Bugaz Spit, bordering the entrance to the Sea of Azov. It consists of several parts the Kiziltash Liman (>Figs. 9.2.1 and > 9.2.2) is retreating differing in morphology because of their geological struc- slowly, whereas the southern part, Vityzaev Spit, borderture, from the southern part of the Mesozoic Scythian ing the same liman is growing because of more abundant Plate in the north to the structures of the GreaLt Caucasus longshore sediment nourishment. in the south (Zenkovich 1958, 1967). The northern end of the Caucasus mountains interEast of the Kerch Strait, for about 20 km unconsoli- sects the coastline at Anapa. This coast has cliffs cut into dated Neogene deposits predominate. The coast is being strongly folded and faulted Cretaceous and seawardintensively eroded and is subjected to landslides. The dipping Neogene flysh (>Fig. 9.2.3). The cliffs between clayey rocks yield little sediment to the beaches, which Anapa and Novorossiiskare up to 200 m high. In the composed predominantly of shelly material. In front of Middle Holocene extensive landslides formed in this area, the cliffs is a wide shore platform. The irregular coastline and they partially survive as small peninsulas. There is results from differential erosion of rock outcrops of vary- little sediment supply from rivers, the continental shelf in ing resistance. Farther southeast, a coastal barrier of quartz the area is narrow and the gradient of the sea floor is steep. sand, about 50 km long, separates a series of large limans Consequently, large waves break on the coast. particularly from the sea. In the west, the barrier is narrow, but it wid- during storms. Below the cliffs are narrow boulder beaches ens to a kilometre in width to the southeast, and carries of flysh rock fragments.

⊡⊡ Fig. 9.2.1 Kiziltash Liman, northern Black Sea, with low terrace formed by a storm surge. (Courtesy Alexei Porotov.)

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⊡⊡ Fig. 9.2.2 The Kiziltash Liman, northern Black Sea from the Taman Peninsula. (Courtesy Alexei Porotov.)

⊡⊡ Fig. 9.2.3 Cliff in steeply dipping Upper Mesozoic flysch sedimentary rocks up to 60 m high Golubaya Bukhta. (Courtesy Ekaterina N. Badyukova.)

The cities of Anapa, Novorossiisk and Gelendjik lie at the heads of bays cut in synclinal folds. From here southward the axis of the Great Caucasus lies only 5–15 km inland. On the 100 km coastline between Anapa and Tuapse the flysch becomes more homogeneous in cliff ­outcrops. There are small bays with narrow beaches of ­fluvial sediment at the mouths of some rivers, and alluvial  clays are exposed on the sea floor at the depths of

between 1 and 8 m. Beaches are poorly developed on headlands, exposed to storms. Seaward from the foot of the cliffs a partially submerged abrasion platform extends down to a depth of 15 m. There are wide beaches and a coastal barrier in Anapa Bay, where coastal defence structures have been built. Farther southeast the Great Sochi resort extends for over 70 km from Tuapse to Adler, and has an alternation of cliffed headlands (>Fig. 9.2.4) and

Russian Black Sea Coast

9. 2

⊡⊡ Fig. 9.2.4 Rugged cliff in Mesozoic limestones and shales near the town of Arkhipo-Osipovka on the northern coast of the Black Sea, with a narrow beach of pebbles and boulders. (Courtesy Vasily A. Dikarev.)

⊡⊡ Fig. 9.2.5 The wide beach of sand and gravel from the Mzymta River at Adler. (Courtesy Geostudies.)

embayments at the mouths of mountain rivers with high sediment yields favoured by the subtropical climate and high precipitation (1,200–1,500 mm/yr). River discharge is highly variable. Narrow shingle and coarse sand beaches occupy coastal embayments, and there is erosion during stormy periods. Longshore drifting is generally southeastward from the mouth of the Tuapse River. Until the middle of the 20th century this sediment flow reached Pitsunda in

Georgia, 150 km to the south east, the longshore drift receiving additional sediment from several other large rivers, such as the Psezuapse, Sochi, Shakhe and Mzymta. In the central part of the Sochi resort the longshore drifting was 30,000 m3 of pebble and 20,000 m3 of sand annually (Zenkovich 1985), and beaches were up to 120 m wide. Southeast of Adler a sandy depositional foreland 8  km long and up to 2 km wide extends between the Mzymta

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Russian Black Sea Coast

and Psou river mouths. It is formed generally of fluvial sand and gravel from the Mzymta River. However, a number of coastal structures, notably groynes, jetties and submarine breakwaters, were constructed in the 1950s and 1960s. They include numerous groynes along the Sochi coastline, and elongated breakwaters at the port of Sochi. A dam was built across the Tuapse River, diminishing its sediment yield to the coast, and in the 1940s–1950s there was extraction of about 100 million m3 of pebbles from beaches for construction purposes. These various measures interrupted longshore drifting and depleted beaches along the coast at and east of Sochi, at Adler (>Fig. 9.2.5) and SE into the Republic of Georgia. Beach nourishment was carried out by local authorities from 1977 onwards to partially restore resort beaches in

the Sochi region, and most of these beaches have survived (Kirillov et al. 1993).

References Kirillov VG, Ivashkov GI, Peshkov VM (1993) The Azov-Black Sea coast of Russia. An experience of artificial beach formation. In: Kos’yan R, Magoon OT (eds) Coastlines of the Black Sea. American Society of Civil Engineers, New York, pp 355–373 Zenkovich VP (1958) Coasts of the Black Sea and the Sea of Azov. (in Russian). Geografizdat, Moscow Zenkovich VP (1967) Processes of Coastal Development. Oliver & Boyd Zenkovich VP (1985) The Eastern Black Sea, USSR. In: Bird ECF, Schwartz M (eds) The World’s Coastline. Van Nostrand Reinhold, pp 473–479

9.3  The Pacific Coast of Russia

Pavel Kaplin · Andrei Selivanov

1.  Introduction The Pacific coastline of Russia, bordering the Seas of Japan and Okhotsk and the Bering Sea, extends for more than 4,500 km (Aseev and Korzhuev 1982; Kaplin 1985). This great length of coastline spans temperate, sub-arctic and arctic climatic zones. Diversity of geological structure and climate conditions result in varied coastal and upper shelf morphology, the main tectonic structures of the Pacific coast region having formed during Mesozoic and Cenozoic times. The mainland and principal islands are mountainous with intermontane depressions and narrow coastal plains (Zenkovich 1967). Mesozoic formations dominate the mainland and Cenozoic formations are prominent on Sakhalin Island, the Kamchatka Peninsula, and part of the Chukotka Peninsula. Block faulting and differential tectonic movement have determined the major features of the coastline. Elevated Pleistocene marine terraces are found only in Kamchatka Peninsula and on the Kuril Islands, and not on the mainland coast, excluding the Chukotka Peninsula. Quaternary volcanic relief occupies large areas of Kam­ chatka Peninsula and the Kuril Islands, but on the mainland there are only a few areas of Mesozoic volcanic rock. On the mainland coast of the Sea of Japan, several embayments formed as volcanic caldera structures. The Late, Pleistocene-Holocene marine transgression invaded former land areas, and subsequent erosional and depositional processes have resulted in the formation of an intricate coastline, including abrasion forms such as fiords and rias, as well as depositional barriers backed by extensive lagoons. Cenozoic block faulting also produced fault scarps, which form sectors of steep coast on parts of the Chukotka and Kamchatka peninsulas. The Sea of Japan is not very stormy, but wave heights reach 2 m in summer and up to 6 m in winter. Ocean swell up to 7 m high and 130–180 m long arrives from the Sea of Japan and the Bering Sea. The Sea of Okhotsk and the Bering Sea are the stormiest seas bordering the Pacific coast of Russia, with prevailing waves 1.5–3 m in height and 20–50 m in length. During 1% of storms in spring and autumn waves reach 12–15 m in height. On the mainland coast of the Japan Sea and the Sea of Okhotsk the

predominant waves arrive from the south, and from the southeast and southwest in the Bering Sea. The northern Bering Sea and northwest Okhotsk Sea coastal areas are ice-covered in winter (for up to 8 months in the north of the area), which limits the effects of wave action on the coasts. However, the influence of floating ice and subaerial denudation processes increases during this season of the year. Tidal influence varies from one sector to another. Microtidal conditions prevail on most open coasts, but on the north of Sakhalin Island, on the north coast of the Sea  of Okhotsk, in several embayments of Kamchatka Peninsula, as well as in some bays bordering the Bering Sea, coasts are mesotidal (2–4 m) and locally macrotidal (as in Penzhina Bay in the north-eastern corner of the Sea of Okhotsk where the tidal range is as high as 13 m). There are extensive intertidal zones, especially in the west of Sakhalin Island, in the Kuril Islands, on the northern coast of the Sea of Okhotsk and on the east of Kamchatka Peninsula. In the Arctic zone thermoerosion, thermodenudation and related processes have strong effects on coastal landforms, and generate large quantities of sediment, including gravel and boulders delivered to rivers and thence to the coast. The gravel beaches of the mainland coast of the Sea of Okhotsk, as well as those of Chukotka and the Kamchatka peninsulas, originated primarily from predominance of various forms of physical weathering in the cold environment.

2.  The Pacific Coastline The southernmost mainland coasts of the Sea of Japan are  characterised by alternation of bedrock outcrops at headlands and bays which are rias (>Fig. 9.3.1), with small depositional sandy beaches in their upper reaches. The climate is rather warm and wet in this area, and chemical weathering of carbonate-rich volcanic and metamorphic rocks is intensive. Several island groups in this part of the sea are similar in morphology. There are coastal lowlands at valley mouths in the Amur Gulf and northward. These lowlands appear to be

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_9.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Pacific Coast of Russia

⊡⊡ Fig. 9.3.1 Posyet Bay, the southernmost part of the Pacific coast of Russian coast. The cape is cut in Cenozoic rocks.

⊡⊡ Fig. 9.3.2 Povorotny Cape (Turning Point) separates a ria coast that is generally transverse to geological structures from a coast parallel to geological structures, with promontories and wide depositional bays.

due to regional and local tectonic emergence. Submerged sandy coastal features at depths of 12–15 m in the western part of the Petra Velikogo (Peter the Great) Gulf, the upper part of the Amur Gulf near the city of Vladivostok, indicate a rapid sea level rise during Holocene. East of Vladivostok, and especially north from Povorotny Cape (Turning Point) (>Fig. 9.3.2), the mainland coast of the Sea of Japan is embayed between bedrock

headlands formed in Mesozoic metamorphic rocks (>Fig. 9.3.3) and intrusive outcrops. The steep underwater slope favours seaward movement of sediment, which remains on the coast only as pocket beaches in embayments. The bays generally become wider northward, and contain extensive depositional features such as sand and shell beaches. Some are partly submerged volcanic calderas, such as Rudnaya Bay. The bays receive large quantities

The Pacific Coast of Russia

9.3

⊡⊡ Fig. 9.3.3 The Two Brothers are stacks on the mainland coast of the Sea of Japan south of Rudnaya Bay. Cliffs cut into Late Cretaceous tuffs, lavas and breccias are cut back during winter storms. There are narrow sand and gravel beaches in embayments, and spits that become tombolos behind residual stacks and islands. (Courtesy E.I. Ignatov.)

of sediment from rivers, which usually have various kinds of spits at their mouths. Several bedrock headlands have been only slightly modified by marine erosion, but others have erosional features such as kekurs (stacks or abrasion residuals in the form of steep-sided rock columns). The west coast of Sakhalin Island is generally erosional, formed on weak Tertiary schists, clays and conglomerates. The height of terraces, mostly erosional, decreases from 60–80 m in the south to 10–20 m in the northern part of Tatar Strait. A former cliff, now inactive, borders 3–5 m Holocene depositional terrace in the south, and passes into an active cliff as the terrace declines northward, The nearshore sea floor has a series of submerged erosion benches and intervening cliffs at depths of 30 m to ­120– 150 m, interrupted by depositional features off several embayments. The east coast of Sakhalin Island has alternations of erosional and depositional landforms. In the south–east are low escarpments in Pleistocene sandy sediment, as in the Terpeniya Gulf, where erosion has been a problem near Vzmorye, and protective structures have failed (>Fig. 9.3.4). There are several rocky Cenozoic outcrops on  the Terpeniya Peninsula which protrudes southward on the eastern side of the Terpeniya Gulf. For over 250 km to the north barriers that separate lagoons and limans from the open sea have been supplied with sand by longshore drifting. The Shmidta Peninsula in the north of the island is one of the few cliffed outcrops of Tertiary rocks.

The Kuril Islands are dominated by volcanic coasts. Some small islands (Matua, Ketoi, Brouton, Chirpoi, Kha­ rimkotan, Atlasova) are recent volcanoes, where andesite lavas and tuffs extend down to the present shores, which have been little modified by marine processes. Other islands are older, inactive volcanoes with coasts on which there has been time for erosional terraces to form. Large kekurs (erosional residuals) are characteristic of these coasts. On the largest islands of the Kuril archipelago (Onekotan, Simushir, Paramushir, Urup, Iturup) there are volcanic slopes and erosional features such as picturesque high cliffs, wave-cut notches, arches and tunnels, with narrow boulder and gravel beaches in embayments. Emerged erosional and depositional terraces occur at heights of up to 100–120 m, and recent depositional features such as spits, tombolos and barriers occur on Kunashir and Iturup and some other islands, but are not widespread. The mainland coast bordering the Sea of Okhotsk is relatively simple in outline. In the south-western corner, Sakhalin Gulf, the coast is dominated by depositional features (spits and barriers) of Holocene age, and the head of  this gulf extensive mudflats are exposed at low tide. A series of Pleistocene depositional marine terraces exists in this area. Farther to the northeast, block folding has resulted in the formation of large islands such as the Shantar Islands, and deep bays such as Akademiya and Tugur. The rock outcrops are rather resistant, and the steep

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The Pacific Coast of Russia

⊡⊡ Fig. 9.3.4 Coastal erosion near Vzmorye in the Terpeniya Gulf, south–east Sakhalin, led to the building of wooden protective structures in 1999 These stabilise the coastline in winter (the picture was taken in November 2000), but are ineffective against late summer and autumn storms. (Courtesy Marina Denisova.)

coasts have been only slightly modified by wave action, but deposition of fluvial sediment has been extensive in Uda Bay. The west coast of the Sea of Okhotsk (Ayanskii Bereg) is fairly straight, following the alignments of tectonic structures and faults. The few embayments have been cut out by marine erosion, and are bordered by cliffs that plunge to depths of several metres. The north coast of the Sea of Okhotsk (Okhotskii Bereg) is characterised by a lobate coastline, transverse to the principal tectonic structures. In embayments, gravel barriers and tombolos enclose coastal lagoons. Most of the barriers formed during the Early and Middle Holocene, and were driven shoreward with a rising sea, which submerged backing lowlands to form lagoons. The largest gravel barrier is more than 200 km, over 2 km wide and up to 8–9 m high, bearing the town of Okhotsk and other settlements. Boulder and gravel beaches dominate in embayments east of Magadan. In Shelikova Bay depositional coastal segments with  gravel and boulder beaches alternate with capes such  as Lisyanskogo and Koni, composed of Mesozoic and Ter­tiary rocks and subject both to wave erosion and ther­modenudation. Extensive intertidal gravel and sand flats occupy the Gizhigina and Penzhina gubas (bays) in the north-eastern corner of the Sea of Okhotsk where the tide range is up to 13 m. The north-western coasts of Kam­chatka Peninsula are dominated by relatively stable cliffs in Mesozoic basalts and andesites, with extensive tidal

flats in the bays. Dissected by deep channels, these sandy and muddy tidal flats are several kilometres wide. South of Ukhtolok Cape and the Khairyuzova River a series of high (6–8 m) Late Pleistocene and Holocene sand and gravel barriers over a kilometre wide extends for more than 600 km, enclosing a chain of lagoons (>Fig. 9.3.5). The barriers have migrated landward, so that Holocene lagoon sediment are found on their seaward slopes at depths of up to 4–5 m. During severe winter storms these barriers may be locally and temporarily breached by waves. The building of solid coastal protection structures has resulted in nearby erosion. In several places the barrier and lagoon chain is interrupted by erosional promontories cut into unconsolidated Pleistocene sands and gravels of problematic origin. On the east coast of Kamchatka the coastline runs transverse to the principal tectonic structures in Mesozoic rocks, resulting in an alternation of uplifted fault-bounded blocks represented by prominent capes such as Shipun, Kronotsk, Kamchatka and Ozerny, and wide arcuate bays. The capes are bordered by wide erosion platforms with residual stacks (kekurs), sometimes with narrow gravel beaches. There are several fiords in the south-eastern part of the coast (Akhomten and Tikhirka bays) and in the north–west (Olyutor Bay), some partly separated from the sea by barriers and loop-like spits fed from both sides. Several parts of the coast have formed along faults, as on Karagin Island, Olyutor Peninsula, with cliffs plunging to 20 m below sea level.

The Pacific Coast of Russia

9.3

⊡⊡ Fig. 9.3.5 The inner (lagoon) shore of a sand and gravel barrier 6–8 m high on the coast of Western Kamchatka, with indications of erosion during high water periods. (Courtesy Serguei L. Nikiforov.)

In the embayments the coast is depositional, with sand  and gravel beaches and barriers enclosing lagoons. On several coastal sectors they form a series of erosional and depositional Pleistocene terraces. Deltas built by ­rivers  have been augmented by sediment drifting from adjacent eroding capes (>Fig. 9.3.6). The beaches contain dark magnetite sands formed from volcanic tuffs. Avacha Bay near the city of Petropavlovsk-Kamchatskii varies in  morphology from a river delta with a wide tidal flat in its head to active cliffs up to 600 m high at its mouth. There are Pleistocene depositional terraces at elevations of up to 60 m in other bays. The northern part of the coast bordering the Bering Sea has been repeatedly glaciated, with several gulfs and bays occupying former glacial depressions. In the northeast of Kamchatka, the Bering Sea borders the Koryack Highland composed of Mesozoic jasper, quartzite and porphyrite, with very picturesque fiords in tectonic depressions. Some of these are characterised by typical sharp saw-like slopes, while others have Pleistocene depositional terraces at their mouths. Large bays such as Korfa Bay are separated from the open sea by extensive gravel barriers and spits (>Fig. 9.3.7). The Komandor (Commander) Islands in the Bering Sea east of Kamchatka are composed of crystalline rocks. The largest of four islands, Bering Island, is generally erosional (>Fig. 9.3.8), with uplifted terraces. Gravel beaches occur only in embayments. The famous fur-seal breeding grounds are here. The extensive Anadyr Gulf has several limans (submerged estuaries partially filled by sediment and separated

⊡⊡ Fig. 9.3.6 South-eastern Kamchatka, the Bering Sea. A river delta formed by deposition of fluvial sediment and augmented by convergent longshore drifting. (Courtesy Alexander N. Apikyans.)

from the sea by barriers), namely the Anadyr Liman, the Kresta (Cross), Provideniya (Providence) bays and others, at its head. Fluvioglacial material fed gravel and boulder barriers which enclose lagoons in former embayments.

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The Pacific Coast of Russia

⊡⊡ Fig. 9.3.7 Chaplina Lagoon in the north-western Kamchatka Peninsula. (Courtesy Alexei V. Porotov.)

⊡⊡ Fig. 9.3.8 Bluffs (formerly cliffs) behind an emerged terrace on Bering Island in the Komandor Islands. (Courtesy Igor S. Branovitskii.)

Some of these barriers are tens of kilometres long, up to 10 m high and over 2 km wide, built up during storms. One of the best examples is the Meechken barrier along the northern shore of Anadyr Gulf. As in Western Kamchatka, most of these barriers have been driven landward, as shown by outcrops of Holocene lagoon peat below sea level on their outer slopes.

The east coast of the Chukotka Peninsula is dominated by fiords in glacial troughs. They are charac­terized by high sinuosity, gentle underwater slopes and shallow thresholds  at the mouths where barriers and loop-like spits have developed (>Fig. 9.3.9). There are rapids over these thresholds, which are usually composed of glacial drift. Occasionally alluvial and marine depositional

The Pacific Coast of Russia

9.3

⊡⊡ Fig. 9.3.9 A looped gravel spit on the coast of Anadyr Gulf. (Courtesy Serguei L. Nikiforov.)

⊡⊡ Fig. 9.3.10 Abrasion residuals on the south-eastern coast of the Chukotka Peninsula. (Courtesy Alexei V. Porotov.)

t­ erraces are found in the heads of the fiords. Abrasion ­residuals are numerous in the bay mouths (>Fig. 9.3.10), and intrusive outcrops form several capes. Dezhneva Cape, the easternmost point of Eurasia, is a bold headland on Mesozoic rocks, and to the north the coast ­bordering the Chukchi Sea marks the beginning of the > Arctic Coast of Russia.

References Aseev AA, Korzhuev SS (eds) (1982) The far east sea coasts of the USSR. (in Russian). Nauka, Moscow, Russia Kaplin P (1985) Pacific USSR. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 863–872 Zenkovich VP (1967) Coasts of the Pacific ocean. (in Russian). Nauka, Moscow, Russia

821

9.4 The Arctic Coast of Russia   Andrei Selivanov*  

1.  Introduction Northern Russia borders the arctic polar seas (Zenkovich 1985). From east to west, these are the Chuckchi, East Siberian, Laptev, Kara, White, and Barents Seas. Their coasts are extremely diverse, and it is possible here only to outline briefly their principal features. Apart from wave action during the summer thaw, the main factors affecting coastal evolution are: (a) shallow seas with strong longshore sediment movement; (b) duration of the ice-free period, which increases from 2.5 months in the Chuckhci and East Siberian seas to 11–12 months in the south-western Barents Sea; (c) influence of various types of sea ice on coasts; (d) high ice content of Pleistocene sedimentary formations on many coastal slopes, resulting in high sensitivity of those slopes to thermodenudation related to slope gradient, aspect, and summer air temperatures; (e) high sediment discharge from the larger rivers and (f) tidal movements, particularly in bays and inlets. Most of the Russian Arctic coast is microtidal, with tide ranges of less than 0.3 m on open coasts increasing to 1.5 m in several bays, but to the west tide ranges attain 11 m in Mezen Bay.

2.  The Arctic Coastline The far NE coast of Russia, bordering the Chuckchi Sea, has a series of Pleistocene depositional terraces of marine, glaciomarine, and glacial origin at elevations of up to 120 m. Capes (such as Yakan, Dezhnev, Shmidt, and Vankarem) are composed of schists and sandstones with granite intrusions. There are fiords at valley mouths. The Koluychinskaya Guba (Guba is a local name for a narrow, deep inlet) in the east of the Chukotka Peninsula is separated from the sea by two wide spits recurved into the bay. The spits have been nourished by strong longshore sediment movement. In the western part of the Chuckchi Sea coasts are various types of lagoons formed during the Late Quaternary (Postglacial) marine transgression, which rose across the sands and gravels of the flat Primorskaya lowland. In Russian literature, a distinction is made between lagoons,

which are marine or estuarine waters behind spits or barriers and limans, which are essentially drowned valley mouths (rias) separated from the sea by coastal barriers. Many lagoons occupy pre-glacial depressions. The enclosing gravelly barriers, up to 7 m above mean sea level, have been supplied with sediment swept in from the sea floor by storm waves and derived from adjacent erosional capes. Most of these barriers are still growing. In general, the Chukotka lagoons have been segmented into numerous rounded lakes as the result of wave regimes related to local winds, as described by Zenkovich (1967). There are also submerged lagoons and barriers on the shelf beneath the Chuckchi Sea. The shallow East Siberian and Laptev seas (prevailing depths less than 50 m) have been shaped during Quaternary oscillations of sea level and resulting migrations of the coastline. Successive low sea level stillstands have left a series of submerged sandy barriers at depths of up to 80 m. The extensive coastal Severo-Sibirskaya (North Siberian) lowland is covered by Pleistocene aleuritic (i.e. coarse silty) sediment with ice inclusions (known locally as the edoma complex). Spurs of mountain ranges of Proterozoic age do not extend to the present coastline. The coastal lowland results partially from an abundant sediment supply from large rivers during floods. The rivers are, from east to west, the Kolyma, Alazeya, Indigirka, Yana, Omolon, and Lena. Their former estuaries have been reduced by alluvial deposition: the Leda is the only river to have built a delta protruding into the open sea. The coastal lowland is generally flat, but with structures such as hummocks and hydrolaccolites of various sizes produced by permafrost, as well as solifluction terraces. There is no definite evidence of morphological features formed during phases of sea level higher than the present. Billings Cape, separating the Chuckchi Sea and the East Siberian Sea, is a depositional foreland consisting of a coastal barrier about 50 km in length and up to 1.5 km wide, backed by an extensive lagoon with transverse sand spits that divide it into a series of rounded lakes over 10 km in width. Billings Cape has formed on the southern coast of De Long Strait, in the wave shadow of Wrangel Island, which is situated about 150 km offshore to the

*Edited version of chapter 9.4 (The Arctic Coast of Russia) in The World’s Coasts: Online (2003) by Andrei O. Selivanov. All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_9.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Arctic Coast of Russia

northeast. The strait is ice-blocked for 9–10 months each year, and waves move sediment shoreward east and west of the ice-blocked area, thereby nourishing the depositional foreland (Solomatin et al. 1998). Chaunskaya Guba (Bay) in the central part of the East Siberian Sea area has been cut into a generally straight coastline. This bay and adjacent coastal segments are bordered by a series of submerged erosional and depositional coastal features at depths of up to 40 m. The bay is partially separated from the open sea by Ayon Island, composed of unconsolidated sediment, and being reduced by thermal erosion. To the west, there are relatively narrow coastal lowlands (up to 4 km) on a 250 km sector where bedrock outcrops form promontories such as Cape Shelagskii, Cape Baranov, and Cape Letyatkin. West of the Kolyma river mouth is a lobate coastline extending over 1,500 km to the large Buor Khaya Bay. Much of this coastline is a low-lying plain bordering a very shallow sea, but there is exposed bedrock at the Svyatoi Nos Cape bordering Dmitrii Laptev Strait between the East Siberian and Laptev Seas. This low-lying coast is submerged by storm surges when northerly (onshore) winds drive the sea up to 30 km inland, depositing muddy sediment on the coast and low hinterland. In contrast, southerly (offshore) winds expose the nearshore sea floor. Tides are here less significant than these displacements. Some sectors of the coast have slopes several metres high, modified by intensive thermal erosion because of melting of buried blocks of ice in the coastal sedimentary formations.

New Siberian Islands separate the Laptev and East Siberian seas. Two of these large islands, Bolshoi Lyakhovskoi and Kotelnyi, are of Pleistocene sediment overlying metamorphic bedrock, which outcrops above sea level on the western side of Kotelnyi Island to form an embayed cliffy coast. However, the greater part of these islands are formed of Pleistocene silty sediment with ice inclusions (edoma complex), subject to thermodenudation (subaerial slope erosion resulting from periglacial processes) and thermoerosion (by waves acting on a permafrost coast) (>Fig.  9.4.1). The islands of Maly Lyakhovskoi, Novaya Sibir, Faddeev, and many others are also composed entirely of Pleistocene deposits with blocks of buried ice (>Fig.  9.4.2). This is a low-lying archipelago on which thermal erosion processes are extremely rapid (Are 1988), the coastline retreating at up to 15 m/year. Various types of ice and snow features may protect the coast from erosion (>Fig. 9.4.3), but several parts of these islands, such as Zemlya Bunge in the central part of Kotelnyi Island, are almost totally overwashed by the sea during storm surges. Most of these islands are likely to disappear in the next few decades (Kaplin 1995). Similar features characterise the SE part of the Laptev Sea coast, with the exception of Svyatoi Nos Cape between the East Siberian and Laptev Seas. Outcrops of pre-Quaternary rocks at elevations of up to 387 m make this cape a unique coastal feature, in contrast with the ice-containing Pleistocene sediment of adjacent slopes, which are degrading rapidly. As the ice melts, the overlying sediment subsides, and there is slumping, collapse, and downwash of ⊡⊡ Fig. 9.4.1 Cliffs on the southern coast of Bolshoi Lyakhovskoi island, New Siberian Islands, are eroding principally by thermodenudation (subaerial slope erosion resulting from periglacial processes) and also by thermoerosion (by waves acting on a permafrost coast). (Courtesy V. Tumskoi.)

The Arctic Coast of Russia

9.4

⊡⊡ Fig. 9.4.2 The coastal slope on NW Novaya Sibir Island, New Siberian Islands contains 80–90% layered ice, and is eroding rapidly. (Courtesy V. Tumskoi.)

⊡⊡ Fig. 9.4.3 Snow patches serve as natural coastal defences against erosion during most of the ice-free period on Zhokhov Island, New Siberian Islands. (Courtesy V. Tumskoi.)

destabilised sediment on coastal slopes. The basal talus is removed by wave action during the summer thaw. Coastline retreat can be as much as high as 50 m/year, as on the small Shelonskie islands, Semenov and Vasilyev (>Fig. 9.4.4). The vast delta of the Lena River is a plain 20–30 m high, much dissected by rivers, particularly in the eastern

and central (frontal) parts. There has been erosion by the thawing of ice-rich frozen ground. In the west, the deltaic coast is bordered by a coastal sand barrier. The finegrained alluvium of the Lena River is being transported mainly to the western part of the delta and distributed over the sea floor. Holocene deposition has culminated in

825

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9.4

The Arctic Coast of Russia

⊡⊡ Fig. 9.4.4 Remnants of the mucheroded Vasilyevski and Semenovski Islands in the eastern part of Yana Bay. They will be consumed by the Laptev Sea in a few decades. West Shelonski Island is in the foreground. (Courtesy V. Tumskoi.)

⊡⊡ Fig. 9.4.5 Rapidly decaying thermokarst mounds in polygonal ice south of the Bykovskii Peninsula and east of the Lena River delta in the north-western part of Buor Khaya Bay, Laptev Sea. (Courtesy V. Tumskoi.)

the growth of a twentieth century delta. The loess-like delta sediment are subject to deflation, and the delta surface is highly variable due to strong winds, storm surges, and periglacial processes. Thermokarst and related processes are typical of the delta margins (>Fig. 9.4.5). West of the Lena River delta, erosional and depositional sectors alternate and the coastline has a generally lobate pattern. Cliffed segments are cut in ice-containing

glacial and glaciomarine sediment of Pleistocene age, the  rate of erosion varying according to the amount of ice  in the frozen ground. Much of the sediment eroded from these cliffs drifts westward along the shore. Farther west, the extensive concavities of Olenek Bay and Anabar Bay have only small river deltas, and generally eroded coastal slopes on promontories. The Khatanga Guba is a large lagoon (liman, a drowned valley mouth enclosed

The Arctic Coast of Russia

by a coastal barrier) in the south-western corner of the Laptev Sea. The Taimyr Peninsula and the Severnaya Zemlya Islands border the Laptev Sea, and have erosional and depositional emerged coastal terraces. The coastline of the large mountainous Taimyr Peninsula is more than 1,000 km long, and highly variable. Fiords and skerry coasts prevail in this area, especially west of Chelyuskin Cape, the northernmost point of Eurasia. The Severnaya Zemlya (Northern Land) Islands are scattered in the sea NW of Taimyr Peninsula. These islands are composed of unconsolidated ice-bearing sediment subject to drastic thermal erosion. Remnants of submerged islands that existed early in the twentieth century are shown as shoals at depths of 3–10 m on modern maps. The islands that persist have an intricate configuration, combining eroding cliffs and depositional features such as sand spits and coastal barriers. West from Taimyr Peninsula and the Pyasina River mouth to Dikson at the entrance to the Yenisei Guba is a smooth coastline about 200 km consisting of cliffs cut in bedrock. The Yenisei Guba and the Ob’Guba farther west are large valley-mouth lagoons (limans) on a low-lying mainly depositional coast, which is subject to submergence by exceptionally high tides and storm surges. The Ob’Guba is the world’s largest liman, about 900 km long and up to 50–60 km wide. The extensive coastline between these bays and the Yamal Peninsula to the west consists entirely of a low plateau in ice-bearing Pleistocene deposits with alternating depositional and erosional segments along the coast. There is rapid thermal erosion on sectors exposed to open sea, notwithstanding the short ice-free period. The average annual rate of retreat of the open coast amounts to 5–10 m. The coasts of the extensive Baidara Bay are fringed primarily by relatively low (up to 20 m) coastal scarps in ice-bearing sediment cut back by wave action and thermoerosion. The present rate of retreat is up to 15 m/year on the west and south facing coasts of the bay (Solomatin et al. 1998). The ice content of these sediment varies from 20 to 30% in sands to at least 50–70% in loams and clays. Prevailing southward longshore sediment movement on the east side of the bay results in bay-head deposition, where there are intertidal mudflats up to 2–3 km wide. A chain of large islands separating the Kara Sea from the Barents Sea are a northward continuation of the Palaeozoic structures of the Ural Mountains. Vaigach Island, with the straits of Yugorskii Shar to the north and Karskie Vorota to the south, is dominated by outcrops of

9.4

Palaeozoic bedrock and beaches and spits of cobbles and pebbles and spits in bays. To the north, the two islands of Novaya Zemlya (New Land) consist of resistant Palaeozoic rocks. In the southeast of Novaya Zemlya, the topography is related to strongly folded structures running parallel to the coast, which is of Dalmatian type. The coastline length of the two islands is over 2,500 km. Glacially smoothed mountains begin on the southern island and extend to the northern island, where they rise to 900 m. Glaciers persist on the northern island, where numerous ice tongues flow down into the heads of fiords, and there is large-scale ice calving on both the eastern and western coasts. The two islands are separated by Matochkin Shar Strait, formed by connecting the heads of two fiords. On the shores of Novaya Zemlya, there are emerged marine terraces and strandflats. The Franz Josef Land archipelago is composed of Jurassic deposits capped by a thick basalt cover. Glaciers have carved deep trenches along depressions and divided the massif into separate islands. The ice-free period on the coast is very short, and shore processes are impeded. Much of the shore is lined with basaltic talus that has not been graded or rounded. The south-eastern part of the Barents Sea, south of Novaya Zemlya, is usually referred to as the Pechora Sea. The coasts are bordered by ice-rich Pleistocene marine and glacio-marine sediment at elevations of up to 120 m. Mezen River opens into tidal Mezen Bay, in the upper part of which the tide range is up to 11 m. Intertidal mudflats (laidas) have a dynamic morpohology, changing during each tidal cycle, and have generally been degraded over recent decades. Morzhovets Island at the entrance to the bay is an example of extremely rapid erosion of icerich coastal scarps. The documented rates are more than 17 m/year, and the island is likely to disappear soon (Selivanov 1996). The White Sea is hydrologically a large tidal inlet from the Barents Sea. Tidal currents and waves enter and leave the White Sea through the wide strait known as The Throat. They transport sediment along the shore, where Intsy Cape is a wide sandy foreland covered with dunes formed at the point where currents converge. The eastern coast of the White Sea is generally composed of Quaternary sediment with isolated outcrops of Devonian and Permian rock. North of the Dvina delta, the coast runs smoothly as far as Cape Voronov. The mouths of the Severnaya Dvina (North Dvina) and Onega rivers are tidal estuaries. Near the mouth of Severnaya Dvina, the largest river flowing to the White Sea, two longshore sediment flows diverge, one

827

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9.4

The Arctic Coast of Russia

­ orthward and the other into Dvina Bay, where a long n spit and Mudyug Island have formed. The outer coasts between the bays, and along the southern White Sea, have eroding slopes in ­unconsolidated Late Pleistocene and early Holocene glacial, glaciolacustrine, and glaciomarine deposits. Pleistocene terraces of marine origin do not exceed 20 m in elevation in this area (Selivanov 1996). The ice content of the sediment is low in comparison with areas to the northeast, and the high proportion of glacial boulders in the slopes and on the beaches prevents rapid retreat. In bays, there are wide intertidal zones of sand and mud (>Fig. 9.4.6), which impede coastal retreat, and some coastal segments have successive coastal barriers, usually Holocene (Selivanov 1996). The only rock outcrop is the small island off the mouth of the Onega River. The pattern of coastline morphology changes drastically on the southern coast of the White Sea near the Vyg River mouth, which is now the outlet from the Belomoro– Baltiyskii Canal (White Sea-Baltic Sea Canal) that has connected these seas for over 60 years. This has had a significant effect in supplying large quantities of sediment to the coastal area. Because of Postglacial isostatic rebound, the coast west of the Vyg River mouth is dominated by erosional slopes in Proterozoic and Early Palaeozoic rocks, extending to the Kola Peninsula. The main coastal feature in this stretch is the Kandalacksha Bay, with typical skerry landscapes. Glacial abrasion forms abound (>Fig. 9.4.7), and relatively narrow boulder and pebble intertidal shores border several bays.

Erosional features become dominant north of Kandalacksha Bay. Typical fiords are rare, being generally filled with glacial deposits that have been subsequently incised by river valleys. Fiard-type coasts prevail. The ­classical series of marine terraces described initially by Lavrova (1960) at elevations of up to 125–140 m and up to 260 m in the center of the peninsula are generally erosional, but depositional remnants contain mollusc shells that have enabled researchers to trace stages in the Postglacial rebound in the area. This terrace series continues northwestward along the Barents Sea coast. In the south-western Barents Sea (The Murman coast), the plateau of the northern Kola Peninsula is composed of Pre-Cambrian gneisses and granites. The latter resist erosion, and the morphology of the Murman coast has been shaped through repeated Pleistocene glaciations. Some straight sectors of plunging cliff run along fault lines, and may have been formed by tectonic uplift along the faults. Coastal erosion has been impeded by the resistance of rock outcrops and by steep nearshore gullies to the depth of some dozen metres, which diminish storm waves. Where the coastal slope is relatively gentle, narrow boulder beaches add to the protection. The western part of the coast is dissected by short and deep fiords. Near the western Russian Arctic border with Norway are the Rybachii Peninsula and Kil’din Island. These are composed of hard metamorphosed Pre-Cambrian rocks, and have cliffs and benches on which wave erosion has been very slow. ⊡⊡ Fig. 9.4.6 A wide sand beach and silty intertidal flat at the head of Onega Bay, SE White Sea. The photograph was taken at low tide in April during the period of ice decay. There is grounded ice on the intertidal flat, but it floats and moves on to the beach during high tide. (Courtesy V. Arkhipov.)

The Arctic Coast of Russia

9.4

⊡⊡ Fig. 9.4.7 In Kandalacksha Bay (White Sea), the shore covered by high tides and storm surges is exposed to rain and other atmospheric processes during low tides. These features were originally formed by glacial abrasion. (Courtesy V.A. Dikarev.)

References Are FE (1988) Thermal abrasion of sea coasts. Polar Geogr Geol 12:1–157 Kaplin PA (1995) The future evolution of the arctic coasts of Russia (in French). Norois 42:37–48 Lavrova MA (1960) Quaternary geology of Kola Peninsula (in Russian). Academy of Sciences USSR Publication, Moscow, Leningrad Selivanov AO (1996) Global sea-level changes during the PleistoceneHolocene and evolution of Sea Coasts. (in Russian). Schwartz, Moscow

Solomatin VI, Sovershaev VA, Mazur II (1998) Dynamics of Arctic Coasts of Russia (in Russian). Moscow University, Moscow Zenkovich VP (1967) Processes of coastal development. Oliver & Boyd, Edinburgh Zenkovich VP (1985) Arctic USSR. In: Bird ECF, Schwartz ML (eds) The  World’s Coastline. Van Nostrand Reinhold, Stroudsburg, PA pp 863–872

9.5  Russian Caspian Coast >11.4

Russian Caspian Coast

9.6  Kaliningrad >8.7

Kaliningrad

829

10.0 Middle East – Editorial Introduction

The Republic of > Georgia is included in the Middle East section because it is difficult to place it anywhere else: it is not part of Europe or of the Russian Federation. The coast runs ESE along the flanks of the Greater Caucasus Mountains, parallel to the strike of Mesozoic sedimentary and metamorphic rocks strongly folded by Alpine earth movements. The Gagra Range, for example, consists of resistant Jurassic and Cretaceous limestones with a very steep slope down to the sea. Beyond the large depositional Pitsunda foreland cliffs in these formations continue up to  Sukhumi, where the mountain front passes inland behind a coastal lowland of Quaternary sediment. The Georgian Trough to the south-east is filled with Tertiary and Quaternary deposits, underlying the valley floor of the Rioni River (Zenkovich 1967). The Black Sea has negligible tides, but variations in water level occur as the result of changes in barometric pressure, wind stress, inputs of rain and river water and seasonal climatic alternations. Wave action is strong ­during storms. Salinity is much lower than in the Mediterra­ nean,  averaging 2.2%, but increasing SW to 3.8% in the Bosporus. Northern  > Turkey is dominated by the Lesser Cauca­ sus, in the form of the North Anatolian Mountains, also of folded sedimentary and metamorphic rocks, and the generally steep coast runs westward along their seaward flank. The Kizilirmak and Yesilirmak Rivers have built substantial deltas here, and the Sinop Peninsula is a protrusion of Pliocene and Pleistocene formations. Further west the steep coast of the North Anatolian (Pontic) Mountains continues, and gives place to the hilly country of the Kocaeli Peninsula, extending past the Sakyara delta. High plateaux border the Bosporus Strait at Istanbul. Western Turkey has mountain ranges and valleys with an E–W structural trend, the ranges running out as promontories and outlying islands while the valleys end in marine gulfs (Sanlaville and Prieur 2005). In the south-west the geological structures swing southward, the Bey Dagliari Mountains running out on the peninsula west of the Antalya alluvial plain. Further east the mountain ranges run SE, then NE in the Toros Mountains,

and the southern coastline follows this trend. Then the mountains curve in behind the wide Seyhan-Ceyhan deltaic plains. East of the Iskenderun Gulf the Amanus Mountains extend southward, curving SW to the Hiznir Cape, and the Orontes valley runs parallel to them on the eastern side. Cyprus is divided into Greek and Turkish provinces (> Cyprus and > Turkey). In northern Cyprus the Kyrenia Range is similar to the Toros Mountains in Turkey, and forms a steep coast. There is a central E–W trough between Mor­phou Bay and Famagusta Bay, and a southern upland, the Troodos Massif, of ultrabasic rocks which influences the steep and rocky western and southern coasts of the island. The east coast of the Mediterranean is hilly on gently folded Cretaceous and Tertiary formations in > Syria, where the coast runs parallel to the N–S anticlinal Jabah Alaouite Mountains. The slopes descend to a low rocky coast. To the south of the Akkar Bay, in > Lebanon a coastal lowland of Quaternary sediment up to 5 km wide stands in front of the Lebanese Mountains, which trend NNE–SSW. There is dune calcarenite along parts of the coastal fringe, and a ridge of Cretaceous limestone comes to the coast at Rosh Haniqra. Terraces on coastal slopes indicate former higher sea levels. In northern > Israel the coastal lowland widens to 5–10 km, but is interrupted where the Mount Carmel ridge comes close to the coast on the Haifa Peninsula. Further south the coastal lowland is up to 20 km wide, and fringed seaward by calcareous sedimentary formations including Pleistocene dune calcarenite. The coastal deposits become more quartzose southward in an area that has received terrigenous sediment from the Nile in Egypt. Mean spring tide range on the east coast of the Mediterranean is small ( Jordan.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Middle East – Editorial Introduction

References Bird ECF, Paskoff R (eds) (1983) Coastal problems in the Mediterranean Sea. University of Bologna, Bologna

Sanlaville P, Prieur A (2005) Asia, Middle East, coastal ecology and ­geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal ­Science. Springer, Dordrecht, the Netherlands, pp 71–83 Zenkovich VP (1967) Processes of coastal development. Oliver and Boyd, Edinburgh

10.1 Turkey

Oguz Erol*

1.  Introduction The coastline of Turkey is 8,333 km long, including about 780 km in European Turkey. The Aegean coastline (3,484 km) is longer and more intricate than the Black Sea coastline (1,701 km) and the southern (Mediterranean) coastline (1,707 km). The various coast types are a direct result of major structural lineaments (>Fig.  10.1.1) that have influenced coastal geomorphology. The mountain ranges are the outcome of folding and faulting that has continued in Quaternary times, and there are frequent earthquakes. The coastline of the Black Sea runs parallel to folds of the high North Anatolian mountain range (Kuzey Anadolu Daglari), and there are almost no coastal indentations or islands. In contrast, the Aegean coastline runs across the structural lines of the Anatolian massif, dominated by E–W horst and graben structures. The coasts of the Sea of Marmara are of a transitional type between those of the Black Sea and the Aegean. The great Anatolian fault, which runs along the northern part of the Anatolian massif, extends along the axis of the very deep Sea of Marmara (greater than 1,300 m), and the E–W coastlines of the Sea of Marmara follow these structural lines, but the N–S coastlines are transverse to structure, and may follow valleys cut along faults, as in the Dardanelles and the Bosporus. Flemming (1978) described the complex plate tectonics, with colliding continental blocks in the Aegean

Sea, which have influenced the west coast of the Anatolian peninsula and the offshore islands. Earthquakes occur from time to time. Structurally, the coastline of the Mediterranean Sea is similar to that of the Black Sea. It runs parallel to the Taurus folded mountain chain, and is also a highland coast, with some elevated terrace plains, but no islands. The northeastern corner of the Mediter­ ranean Sea is also structurally controlled, including the plains of the Qukurova and the graben-like Gulf of Iskenderun. There are terraces at various levels, indicating that the sea has stood at higher levels relative to the land in Quaternary times and that there have been tectonic displacements. The Black Sea was a freshwater lake in Late Pleistocene times, and is thought to have been at a lower level than the Mediterranean Sea until water flooded in over a threshold in the Bosporus about 6,000 years bp. This would explain the discoveries of archaeological evidence of man-made structures out on the continental shelf, for example a house dated 7,500 years bp at a depth of 95 m below sea level 20 km east of Sinope. Soon afterwards, the coast was explored by ancient Greeks, and has associations with Jason and the Argonauts. The coast of Turkey has a Mediterranean climate, with a mild rainy winter and a hot dry summer. Istanbul has a mean monthly temperature of 5°C in January and 23°C in July and an average annual rainfall of 720 mm; Izmir has 8°C

⊡⊡ Fig. 10.1.1 Major structural lineaments as they affect the coast of Turkey.

*Edited version of chapter 10.1 (Turkey) in The World’s Coasts: Online (2003) by Oguz Erol (University of Istanbul, Turkey). All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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10.1    

Turkey

in January and 27°C in July, with an average annual rainfall of 700 mm. However, the Black Sea coast has rainfall all year. Most of the rivers rise in the interior uplands. Some (e.g. the Kizilirmak) have cut transverse valleys through mountain ranges to reach the Black Sea and Mediterranean coasts, others (e.g. the Kucuk and Büyük Menderes) flow into grabens opening to the Aegean coast. Westerly winds predominate, and waves arrive mainly from the NW and NE on the Black Sea coast and the SW on the Mediterranean coast, but the Aegean coast has a shorter fetch and many islands, and wave action is generally weaker. Tide ranges reach only a few centimetres, but meteorological effects such as pressure fluctuations, onshore winds, and storm surges can raise sea level by up to 3 m. Salinity is about 20–25 parts per thousand, somewhat lower than the Mediterranean Sea, but increasing in the southwest to around 38 parts per thousand in the Bosporus. Rivers draining into the Black Sea have poorly developed estuaries because they descend swiftly to a microtidal sea of relatively low salinity. Most end in deltas rather than estuaries.

country, cut partly through Ordovician, Devonian, and Carbon­iferous rocks, and opens northward into the Black Sea. It has an irregular bottom profile with some rocky protrusions, and if it originated as a river valley there has been subsequent tectonic disruption. The Black Sea coast of European Turkey is relatively straight and partly cliffed WNW past Agacli, with the Terkos Golu lagoon enclosed by a sandy barrier across a former embayment. The coast has wide sandy beaches and extensive dunes, which become higher to the west. Cliffs and bluffs then border a 150–250 m coastal plateau on Pliocene formations, backed by the Istranca mountains as the coast curves northward past Midye. The bay at Igneada is partly sheltered by a promontory to the northeast, and about 16 km to the north the Bulgarian border crosses the coast. The coast has wide sandy beaches and extensive dunes, which become extremely high to the west, beyond which the Istranca mountain chain forms a cliff coast. Cliffs are also developed along the edge of the 150–250 m Plioce.

2.  The Coast of European Turkey

The eastern part of the Black Sea coast is dominated by the nearby high mountains, consisting of Mesozoic-Tertiary igneous and sedimentary formations. In some sectors, forested mountains descend directly to the coast, which has a succession of cliffy promontories alternating with bayhead beaches, mainly of shingle (Buachidze 1974). Heavy mineral concentrations are a common component of the beaches and inner shelf (Evans 1971). The steep coast includes Upper and Middle Pleistocene terraces 5, 10–15, and 20–40 m above sea level, a high terrace bearing fluvial deposits at 60–70 m and Lower Pleistocene to late Pliocene surfaces, partly erosional and partly depositional up to 200–250 m. The Rioni (Coruh) River flows northward through a deep gorge at Artvin and across the Georgian border to a delta just south of Batum. Sediment carried down this river has been deposited to form the delta, but in recent decades there has been erosion of the delta shores, and this will intensify if the supply is curtailed by dam construction in Turkey. From the Georgian border to Trabzon (Trebizond), the coast is generally steep, with bold rocky promontories and stacks (>Fig. 10.1.2) and sectors of narrow alluvial plain fringed by gravelly beaches, particularly around the mouths of rivers. Longshore drifting, mainly from west to east, widens beaches on the western side of harbour breakwaters, as at Rize, leaving those to the east narrow and

From the Greek border along the Meric River, the southfacing coast with cliffs and bluffs in Miocene rocks follows a fault line along the edge of the Saros Gulf (Saros Körfezi) past Ibrice. The gulf shallows eastward to a deltaic­ lowland, and on the southern side is the steep coast of the Gelibolu (Gallipoli) Peninsula, with high ridges in Eocene to Miocene rocks. The long southern coastline of the Saros graben may be an extension of the greater Anatolian fault, which extends eastward to the Sea of Marmara and across northern Anatolia. The coast declines to Ilyasbaba Point, on the northern side of the entrance to the steep-sided Canakkale Bogazi (Dardanelles), with the high island of Gökceada to the west. The Dardanelles connect the Aegean Sea to the Sea of Marmara, and originated as early Quaternary river valleys established along fault lines: they have been repeatedly invaded by the sea during Pleistocene interglacials and the Late Quaternary marine transgression. The steep coast continues past Gelibolu (Gallipoli) to Tekirdag, and on Miocene formations east to Bakirköy and Istanbul. It is backed by a Neogene plateau with ­elevations varying from 200 to 250 m, and rivers entering the sea deliver sand to form spits that enclose small lagoons west of Yesilköy. The narrow, slightly winding Bosporus channel (Karadeniz Bogazi) is incised into hilly

3.  The Black Sea Coast

Turkey  

   10.1

⊡⊡ Fig. 10.1.2 Cliffs and stacks on the Black Sea coast near Trebizond. (Courtesy Geostudies.)

depleted. The coastal highway follows low terraces between 5 and 20  m above sea level, and tunnels through spurs. Many sectors of coast have been armoured with boulders to protect the highway, and some low terraces have been widened artificially. Near Pazar, there are low eroding cliffs in weathered volcanic rock (>Fig. 10.1.3) and segments of almost flat shore platform produced by weathering, with residual knobbly protrusions (embedded volcanic bombs). Some of the deeply incised valleys are bordered by high river cliffs exposing sand and gravel terrace deposits, as at Ardesen. At Giresun, T-shaped groynes have been built in an attempt to trap and retain a protective shingle beach, but with little success. West of Ordu, the coast curves out past Persembe to  Cape Cam, on a promontory of Cretaceous rocks (>Fig. 10.1.4). There are cliffed spurs of dark grey volcanic rock, agglomerate, and tuff, with some stacks and islets, and planed-off shore outcrops, and intervening shingle coves. A steep wooded coast then runs west to Yasun Burnu, and swings southwest to the mouth of Balaman River. The mountains are fronted by a narrow coastal plain extending past Unye, where there are stacks of volcanic rock standing in the sea. Thgen the coastal plain widens toward the large delta of the Yesilirmak River. Near Terme is the Simenlik Golu lagoon on the northeastern shore of the delta, the river flowing to the sea at Civa Point where the coast is prograding (Erol 1983). The delta shores are sandy beaches, backed by marshes and small lagoons. On the landward side, the delta is backed by hilly country, interrupted by the broad, terraced Yesilirmak valley.

On the western flank, the delta shore curves round to Samsun, where a steep slope descends to a narrow coastal plain. This widens near the mouth of the Murat River, and to the north west the coastal plain again widens, this time to the large prograding Kizilirmak delta (Erol 1983). This has the large Balik Gölü lagoon on its eastern coast, flanked by a barrier spit that has grown southward. The delta shore is again sandy, backed by reedy and rushy marshy areas, and the river mouth opens to the apex of the delta at Bafra Point, constricted by spits. The Kizilirmak River carries a load of sand and small pebbles (>Fig. 10.1.5), delivered to beaches on the delta shore. The continental shelf is narrow along the eastern part of the Black Sea coast, but widens to between 25 and 50 km off the Yesilirmak and Kizilirmak deltas, forming submarine lobes with an asymmetrical eastward trend, probably caused by dominant easterly currents. There are the remains of submarine valleys off the two deltas. The delta narrows to westward to Alaçam, where the coast is steep and partly cliffed behind a wide continental shelf. The Sinop peninsula is set forward from the high North Anatolian Mountains and consists of Upper Plio­ cene and early Pleistocene sedimentary rocks, bordered by fairly high cliffs. The rocks have yielded Uzunlarian and Karangat fossils of late Pleistocene age. From the northwest corner of the Sinop Peninsula at Ince Burun, the coast swings south and west to Ayancik. The mountain ranges continue westward, consisting of Palaeozoic to Tertiary sedimentary rocks with associated volcanic formations. The coast is dominated by high cliffs

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⊡⊡ Fig. 10.1.3 Erosion west of Pazar. (Courtesy Geostudies.)

⊡⊡ Fig. 10.1.4 Steep coast at Cape Cam. (Courtesy Geostudies.)

with small bay-head shingle beaches and there is an intermittent narrow coastal plain. To the west, step-like foothills and segments of early Pleistocene plateau occur. In this western part of the Black Sea coast, there is an increase in northeasterly wind and wave action, producing some westward drift of beach material, but this is often exceeded by the effects of occasional northwesterly storms. The headland at Cape Baba shelters a bay at Eregli (Heraclea) and narrow coastal plains front steep ranges westward past Kocaali, broadening to the Sakarya delta.

This is a blunt delta extending about 40 km along the coast on either side of the river mouth, where longshore drifting from west to east has built a small spit. Alluvial sediment has been distributed by waves and by the easterly surface currents and westerly subsurface (20–30 m) deeper ­counter-currents. The evolution of the Sakarya delta raises problems. The river carries a silty load, but the coast for at least 10 km on either side has prograded sandy beaches and dune ridges (>Fig.  10.1.6). These were probably derived from the yellow Eocene (Lutetian)

Turkey  

   10.1

⊡⊡ Fig. 10.1.5 Sand and pebble load of the Kizilirmark River. (Courtesy Geostudies.)

⊡⊡ Fig. 10.1.6 Prograded beach west of the Sakarya delta. (Courtesy Geostudies.)

sandstones, but their origin and mode of delivery require more research. The coast runs out to a lobate promontory of Cretaceous formations at Kefken Adasi, and to the west, past Sile, is dominated by the edge of a high plateau. Bayhead beaches alternate with cliffed promontories and volcanic rocky headlands, and the plateau is incised by a number of small valleys opening seaward, and then the northern mouth of the Bosporus.

The Sea of Marmara is very deep, and is a subsided basin within the greater Anatolian fault system. Its coasts are complex and structurally controlled with a series of small gulfs or embayments and islands. High, steep coasts predominate, but small deltas are common at the mouths of incised valleys. From the southern end of the Bosporus, the coast runs southeast and alongside the long Izmit Gulf, where there has been extensive urban and industrial development.

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Turkey

Izmit Gulf occupies a graben, which runs eastward, and is occupied by Sapanca Lake, about 50 m above sea level. This may formerly have been part of a higher Izmit Gulf. To the south is the hilly and mountainous peninsula that runs out to Boz Burun, a headland protruding into the Sea of Marmara. To the south is the smaller Gemlik Gulf, on the southern side of which is a narrow coastal plateau, across which flows the Simav River. To the west, beyond Badirma is a narrow isthmus linking a granitic promontory that rises to a peak 803 m above sea level. Marmara is a similar outlying high island. Erdek Bay, to the south, has a sandy coast fringing a narrow coastal plain across which flow the Gönen and Kocabas Rivers. The coast swings north to Ince Point, then west along the gulf that narrows to the steep-sided Dardanelles.

4.  The Aegean Coastline The coastline of the Aegean Sea includes structurally controlled indentations, with many fault-bounded graben-like features forming bays between uplifted horsts that form promontories. It has steep scrub-covered slopes descending to alternations of rocky cliffs and deltaic lowlands with beaches and muddy areas in sheltered bays. Along the Aegean coast surfaces and terraces have been identified at 150–200, 90–110, 30–40, and 16–18 m above sea level (Eisma 1978; Kraft et al. 1980). Submer­ gence during the Holocene epoch was followed by progradation of river deltas such as the Gediz, the Havran, the Kucuk Menderes, and the Buyuk Menderes. Sea floor contours in the Aegean surrounding the promontories and islands suggest a complex tectonic history, which may have included the dislocation and submergence of early Quaternary or pre-Quaternary plateaux and sediment-filled valleys. Mehmetcik Burnu is the point on the southern side of the western mouth of the Dardanelles, close to Kumcale, where there are fluvial and marine terraces, and the Kara Menderes valley is the site of ancient Troy. The coast runs southward behind Bozcaaada island, and the volcanic rocks of the Biga Peninsula come to Baba Point. To the east are cliffs following the lines of faults along the northern coast of Edremit Gulf, which is fed by the Havran River and sheltered by the Greek island of Lesbos. The coastline becomes intricate around Ayvalik and runs out to the hilly Candarli promontory. Candarli Gulf receives the Bakircay River, and the coast is irregular along to Foca. The Gediz River, draining another E-W structural trough, comes to the coast on the northern side of the wide Menemen plain, bordering the angled Gulf of Izmir. This

lies east of the mountainous peninsula of Karaburun, which has steep scrubby coasts with beaches in intricate bays and small headlands. East of Doganbey Point, a broad bay receives the Kucuk Menderes River near ancient Ephesus (Efes), and the coast curves out along the mountainous Samsun Peninsula. North of the broad Buyuk Menderes valley, which drains another large graben. This was formerly a major marine embayment, on the shores of which were ancient Greek and Roman cities such as Priene, Miletus, and Heraclea, but it has been filled by prograding alluvium. Progradation continues on the sandy deltaic coast, where spits have been built by longshore drift enclosing lagoons that are rapidly silting. Gulluk Bay has numerous elongated inlets between long promontories, which continue round the Kefaloka Peninsula. To the south, another deep re-entrant, Kerme (Gokova) Gulf, is a fault-bounded graben that runs back to Iskele, and has received very little sediment infill because it lacks a major drainage system in the hinterland. On its southern side is the mountainous Marmaris Peninsula, with long, narrow ridges running out as the Resadiye and Daracya peninsulas. The coastline is very intricate, with alternating cliffed headlands and bay-head beaches, and there are many small islands. The Daracya peninsula runs south to Kara Point, looking out to the Greek island of Rhodos, and the southern coast of Turkey begins.

5. The Southern (Mediterranean) Coastline The coastline of Mediterranean Turkey is similar to that of the Black Sea in that it is longitudinally parallel to the Taurus Mountains, consisting of Palaeozoic to Tertiary folded formations with much karstic limestone topography. The only major exceptions are north-south structural features in the embayment at Antalya, bordered by high cliffs and plateaus of travertine, and the northeastern embayment, dominated by the Qukurova deltaic plain and the tectonically downwarped Gulf of Iskenderun. Much of the coast is backed by steep scrub-covered slopes. East of the Daracya Peninsula, the coastline consists of many irregular bays and promontories extending past the bay, which receives the Dalaman Kirenis River to the narrow-necked Mesozoic peninsula of Kutoglu. Beyond is Fethiye Bay, with an irregular coast curving down to Yedi Point. Koca River flows southward in a valley that ends in a blunt delta, wave energy having increased along the coast east of the protection of Rhodos. The continental shelf is narrow here, and along the south-facing coast with

Turkey  

its many small steep-sided promontories to Gelidonya Point. The steep coast swings northward on the western side of Antalya Bay, at the head of which is a broad alluvial lowland. There are cliffs of travertine at Antalya, and the coast east to Alanya is generally sandy, with outcrops of beach rock, and dunes in ridges running parallel to the coastline. Fluviomarine terraces extend along the coastal lowland, with travertine terraces of Pleistocene age at 300–260, 190–220, and 60–70 m. There are also river terraces 140–120, 100–82, 72–60, 58–72, 24, 18, 10, 6, and 2 m above sea level. East of Alanya, the Taurus Mountains come to the coast, and there are high steep and cliffed coasts, mainly on Mesozoic limestones and Palaeozoic schists. Sandy beaches fringe small deltaic plains at the mouths of rivers, but the coastline is dominated by cobble and boulder beaches at the foot of the cliffs. The broadly lobate coast reaches a southernmost point at Anamur Burnu, which faces across the 70 km strait to the north coast of Cyprus. The cliffs come to an end at Tasucu, beside the lobed delta of the Goksu River. The delta has a spit that extends southwest to Bagase Point, shaped by easterly wave action on a coast that is somewhat protected from southwest waves by the island of Cyprus. It incorporates several small lagoons. Northeast of the delta, from Silifke to Mersin, the lower slopes of the Daglari Mountains descend to the shore on a coast that faces southeast. The continental shelf widens eastward in front of the large Seyhan–Ceyhan delta. The extensive Cukurova alluvial plain has been built by the two large rivers, the Seyhan to the west and the Ceyhan to the east. Much of the deltaic plain was formerly swampy, but has been drained for agriculture. Southeasterly waves have formed broad beach ridge plains along the coastal fringe, enclosing some lagoons, the largest of which is the Akyatan Lagoon. Evans (1971) described three sets of Pleistocene to Holocene beach, dune, and lagoon systems in the delta of the Seyhan River. The Ceyhan delta is bordered by a series of terraces that occur along the northern side of the Holocene delta plain at 10–15, 20–35, 50–60 m. The Ceyhan River has cut

   10.1

through a ridge of Miocene rocks, and has changed its lower course from west to east, at present flowing out northeast behind a spit into a bay beside the Gulf of Iskenderun. The Gulf of Iskenderun is rectilinear and structurally controlled, between the Misis Mountains to the north and the Nur Mountains to the south. There are small beaches to the northeast and large depositional fans extending into the gulf along the foot of the Nur Mountains. The mountainous peninsula of Triassic and Cretaceoius rocks runs out to Hinzir Point, and the coast swings southeast to the mouth of the Orontes River. This flows across the broad inland Amik agricultural plain, with includes the swampy depression occupied by the Amik Lake, and out through a narrow defile at Antakya (ancient Antioch) to the delta. Near the river mouth are the remains of the ancient Greek and Roman harbour of Seleucia ad Orontes, which stands above present sea level. A stairway of marine terraces descends from 110–140 m with stages 83–100, 58–75, 35–50, 15–20, 7, and 2 m above sea level, and episodes of Holocene emergence are indicated by solution notches at 2.5, 1.4, and 0.5 m in cliffs at Cevlik, just north of the delta.

References Buachidze IM (1974) Black Sea shelf and littoral zone. In: Degens ET, Ross EA (eds) The Black Sea. American Association of Petroleum Geologists, Tulsa, OK, pp 308–316 Eisma D (1978) Stream deposition and erosion by the eastern shore of the Aegean. In: Brice WC (ed) Environmental history of the Near and Middle East. Academic Press, New York, pp 67–79 Erol O (1983) Historical changes on the coastline of Turkey. In: Bird ECF, Fabbri P (eds) Coastal problems in the Mediterranean Sea. Inter­ national geographical union, commission on the coastal environment, Bologna, pp 95–108 Evans G (1971) The recent sedimentation of Turkey and the adjacent Mediterranean and Black Seas: a review. In: Campbell AS (ed) Geology and history of Turkey. International Publications Service, New York, pp 385–406 Flemming N (1978) Holocene eustatic changes and coastal tectonics in the northeast Mediterranean: implications for models of crustal consumption. Philos Trans R S Lond A, 289:405–458 Kraft JC, Kayan I, Erol O (1980) Geomorphic reconstructions in the environs of Ancient Troy. Science 209:776–782

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10.2 Cyprus

Yaacov Nir

1.  Introduction Cyprus, the third largest island of the Mediterranean (9,250 km2), is located in the north-eastern part of the Levantine Basin, about 85 km south of the Turkish coast and some 100 km west of Latakiya on the Syrian coast. It has a coastline about 776 km long. The northern part of the Republic of Cyprus has since 1975 been occupied by Turkey, extending from the southern side of Morphou Bay eastward to the Karpas Peninsula, then down the east coast to a point south of Famagusta. The island has many prominent capes, mainly of limestone and calcarenite, between bays (Thrower 1960). The importance of the capes was recognised in antiquity by the name Ceratis, the horned island. There is a steep-sided northern range of limestone and marble, the Kyrenian Range, running W–E with peaks rising to around 900 m and passing eastward into the elongated Karpas Peninsula. This is bordered southward by the Plain of Mesaoria, a gently undulating lowland that runs from Morphou Bay in the west to Famagusta Bay in the east, and then the Troodos Mountains, with peaks rising to more than 1,950 m (Mount Olimbos) in the south-west and deeply incised radial valleys. There is evidence of differential uplift during the Quaternary, with marine terraces at various levels. Cyprus has a Mediterranean climate, with a mild rainy winter and a hot dry summer: in the Troodos Mountains annual rainfall attains 1,200 mm, but in the Plain of Mesaoria it diminishes to about 300 mm. Most of the rivers rise in the Troodos Mountains, flowing as torrents in winter but becoming dry watercourses in summer. Westerly winds predominate, and the west coast is stormier than the east. Tides are semidiurnal, with maximum ranges of about 30 cm, but meteorological effects such as pressure fluctuations, onshore and offshore winds. Four distinct morphological features characterise the morphology of the island, which has emerged since the Pliocene: the Kyrenia Arch, the Mesaoria Plain, the Troodos Mountains and the Mamonia Hills. These are related to the pattern of geological outcrops.

The Karpasia Peninsula, the elongated narrow peninsula in the NE side of the island, represents a section of the Alpine Orogeny chain of southern Anatolia. It is a quite narrow, somewhat crescentic arch-shaped ridge, running for about 150 km along the northern part of the island, very close and parallel with the coastline. A narrow plain fronts the foothills of the Kyrenia range. Numerous very short streams are typical of the northern slopes of this ridge, gushing into the sea during rainy days. Its southern flanks are also steep, drained by the Mesaoria Plain’s rivers to Morfou Gulf in the west and to Famagusta Gulf in the east. The Kyrenia Terrane is mostly metamorphic represented by rocks from the Permo-Carboniferous Period to  the Neogene. The main rocks are metamorphosed ­limestones, dolomitic limestones, chalks, marls, cherts, and greywackes. The Mesaoria Plain spans the island from Morfou Gulf in the west to Famagusta Gulf in the east, some 90-km long, and between 25 and 40 km in breadth. It has a smooth topography and is characterised by small number of relatively long rivers: three pour their water to the Morfou Gulf, while two originating in the Troodos in the south drain the central and eastern plain to the Famagusta Gulf. The lithology of the plain comprises of Pliocene to Holocene sediment, mostly including conglomerates, gravels and marls. The Troodos Range is an igneous massif, with a core of plutonic rocks surrounded by a zone of intrusives with an outer fringe of pillow lavas along its foothills. The Troodos Mountain range which reach 1,952 m at its highest peak of Mount Olympus, has the highest precipitation rates of the island and is therefore the most important source of highly torrential rivers that transport mixture of fine and coarse sediment to the lowland and the sea. The Mamonia complex of Hilly South Western Cyprus, characterised by the cherty-shale, sandstone-limestone and chalky sequences. The complex forms the foothills of the high Troodos Mountains, gradually approaching the coast. In a few regions, the hilly terrain reaches the sea to form relatively steep and high cliffs.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Cyprus coast has numerous beach compartments, separated by headlands. Many of the beaches are gravelly. Predominant longshore drifting is from west to east on the northern and southern coasts of the island and southward on the western and eastern coasts.

2.  The Coastline of Cyprus The coastline is described in an anticlockwise sequence, starting and finishing at Cape Arnauti, in the north-­ western corner. From Cape Arnauti to Paphos (50 km) the west coast of Cyprus is exposed to a long westerly fetch, and has strong wave action and longshore currents. From Cape Arnauti southward there are sectors of cliffed and rocky shore (as on the Akamas Peninsula) separated by sandy beaches, some with exposures of beach rock. On the SW coast between Paphos and Cape Aspro (50 km) there are cliffy outcrops of calcarenite and associated beaches of brown calcareous sand. Rivers draining the high Troodos Mountains deliver loads of sand and gravel to the coast. There are promontories related to WSW–ENE faults and coves containing beaches of sand and gravel. Southeast of Paphos is the limestone stack known as Aphrodite’s Rock (>Fig. 10.2.1). Between Curium and Cape Zevgari (15 km), Episkopi Bay has beaches of well-rounded pebbles supplied by the Kouris River. These diminish in size along the coastline between the river mouth and Cape Zevgari and also

­ orthward, the beach becoming sandy in Curium Bay. n They have been extensively quarried. From Cape Zevgari to Cape Gata (20 km) is the southern coast of the Akrotir peninsula, a major double tombolo. There are coastal dunes on the Akrotiri Peninsula. The lake on the peninsula was shown as a bay on a sixteenth century map (Ortelius 1573), and has been separated from the sea by recent deposition. The west coast has cliffs up to 30-m high. The southern coast is a narrow ridge of Plio-Pleistocene formation with a cliffed and rocky shore and sporadic narrow beaches. On the eastern coast is a sandy beach that becomes gravelly northward towards Limassol. Cape Gata to Cape Kiti (75 km) is a curved coastline extending from Akrotiori Bay eastward. The Limassol coast is strongly urbanised, with many hotel and marinas. Sea walls were built to stabilise the coastline, but the beach has been depleted, and nearshore breakwaters have failed to trap much sand. To the east, the coast is low-lying bordering a coastal plain, but there is an emerged beach 3–12 m above sea level at Mari (Pantazis 1965). Rivers draining from hinterland ranges have supplied gravel to beaches, and accretion on the western sides of groynes indicates eastward longshore drifting. Near Moni sectors of chalky cliff with chert horizons are fronted by narrow beaches, typically with gravel beach cusps (>Fig. 10.2.2). Between Cape Kiti and Cape Pyla (40 km), the beaches of Larnaca Bay are in general narrow, consisting of sand and some pebbles derived from shore outcrops.

⊡⊡ Fig. 10.2.1 The stack at Petra tou Roumiou, legendary birthplace of Aphrodite.

Cyprus  

   10.2

⊡⊡ Fig. 10.2.2 Governor’s Beach, between Zyyi and Limassol beach, consists of shingle behind sand, and shows beach cusps.

⊡⊡ Fig. 10.2.3 A gully cut in calcarenite near Ayia Thekla.

These become a barrier enclosing brackish lagoons south of Larnaca. There are sectors of marine terrace about 2 m above sea level, composed of well-rounded small pebbles. The Pyla peninsula coast is barren and rocky. From Cape Pyla to Cape Greco (25 km) is a south-facing embayment. East of Cape Pyla former dunes have disappeared because of quarrying or housing ­development. Low cliffs of calcarenite are extensive, incised by elongated gullies (>Fig. 10.2.3). There are several sandy embayments between low rocky headlands. White sandy beaches near Ayia Napa are composed of foraminiferal sand. The cliffs become higher eastward to Cape Greco, an outcrop of Lower Miocene limestone. At Cape Greco, the coast turns NNW, and between Cape Greco and Cape Elea (60 km) is Famagusta Bay. This east-facing coast is dominated by low calcarenite

cliffs with an irregular outline and basal notches (>Fig.  10.2.4). The shore is often gravelly or bouldery. Sandy beaches occupy some embayments, and there are exposures of beach rock. Famagusta formerly had a wide sandy beach, but quarrying has depleted it (>Fig. 10.2.5). North of Fama­gusta, the Yialias and Pedhieos Rivers flow to the sea, but their sediment load is mainly fine and  they have not built deltas or extensive beaches. However, they have contributed to deposition that has stranded the ancient port of Alasia up to 3.5 km inland. Further north, there are submerged calcarenite ridges extending up to 700 m offshore have been planed off by the sea. Cliffs in calcarenite overlying Plio-Pleistocene chalky rocks run out to Cape Elea. From Cape Elea to Cape Apostolos Andreas (70 km) is the south-eastern coast of the Karpasian Peninsula At

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⊡⊡ Fig. 10.2.4 A low cliff in dune calcarenite with a basal notch at Farkonia Beach on the east coast of Cyprus.

⊡⊡ Fig. 10.2.5 Depleted beach north of Famagusta.

Cape Elea the coast turns northward for a few kilometers and then continues in its northeastern trend. At Moulos, a  rocky ridge parallel to the coastline is attached to the mainland by a broad tombolo (>Fig. 10.2.6). Chalk cliffs extend round headlands and behind inset embayments, bordered by shore platforms that truncate the dippng rocks and carry a veneer of sand (>Fig. 10.2.7). The sandy beaches are generally narrow, except at Ayios Simon and

Koma tou Yialou. They are backed by dunes locally, as at the ­Pak­hyammos tombolo (>Fig. 10.2.8) where a small rounded islet has been linked to the mainland by a sandy isthmus. In places, dunes have spilled inland up to a few hundred metres. Longshore drifting is eastward on this part of the coast. A small group of islands, the Kleides Islands, lie off Cape Apostolos Andreas, continuing the structural trend of the Karpasian Range.

Cyprus  

   10.2

⊡⊡ Fig. 10.2.6 The broad tombolo at Moulos on the south coast of the Karpasian Peninsula. (Courtesy The geological Survey of Cyprus.)

⊡⊡ Fig. 10.2.7 Shore platform cut in chalk near Rhizokarpaso on the south coast of the Karpasian Peninsula. (Courtesy The geological Survey of Cyprus.)

From Cape Apostolos Andreas to Cape Kormakitis (185 km) the north coast of the Karpasian Peninsula is fringed by narrow beaches of calcareous sand derived from the erosion of local outcrops. At Ronnas Bay, dunes have spilled inland, climbing into the Karpasian foothills. Beach rock is exposed locally. To the west, the coast runs along the base of the steep Kyrenian Range, mainly of metamorphic rocks. Short, steep rivers descend to small

embayments. West of Ayios Amvrosos these rivers have built small deltas. There are cliffs cut in chalk and Pleistocene calcarenites, extending westward past Kyrenia to Cape Kormakitis. From Cape Kormakitis to Karavostasi (35 km) is the coast of Morfu Bay. The western side of Cape Kormakitis is irregular, with cliffy headlands and small embayments cut in Pleistocene calcarenites, limestones and marls.

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Cyprus

⊡⊡ Fig. 10.2.8 Air photograph of the tombolo at Pakhyammos. (Courtesy The geological Survey of Cyprus.)

⊡⊡ Fig. 10.2.9 Dunes spilling inland at Ayia Irini on the east coast of Morfu Bay.  (Courtesy The geological Survey of Cyprus.)

The coast then continues southward beside Morfou Bay at the western end of the Mesaoria Plain and has a smooth outline, cut in soft Plio-Quaternary sands, silts, clays and gravels. The sandy beach is backed by dunes of mixed quartzose and calcareous sand that are spilling inland at Ayia Irini (>Fig. 10.2.9), and otherwise covered by pine forests. In the south of Morfu Bay, the coast curves westward, and rivers from the Troodos Mountains have supplied pebbles to beaches (>Fig. 10.2.10). These are being carried along the shore by northward longshore drifting. Lines of sand dunes about 800 m inland from the southern coast of Morfou Bay mark former coastlines associated with a 6 m emerged beach (Wilson 1956). From Karavostasi to Pomos Point (28 km), a steep coast borders the Tillyrian foothills NW of the Troodos Mountains, developed mainly on volcanic formations. There are cliffs and rocky shores, and beaches of volcanic rock pebbles and sand brought down by rivers. Khrysokhou Bay lies between Pomos Point and Cape Arnaoutis (36 km). At Pomos Point, a cliffed coast with rocky shores and cove beaches (>Fig. 10.2.11) declines to a lowland behind Khrysokhou Bay and Khrysokhou River delivers sand, silt and gravels to the shores of this bay. It is exposed to westerly wave action, which generates eastward longshore drifting along the bay shores. The beaches are narrow and gravelly south from Pomos, but widen in the southern part of the bay. West of Latsi the coast consists of low cliffs bordering the Akamas Peninsula, with sand and pebble beaches in coves and embayments sheltered from westerly storms.

Cyprus  

   10.2

⊡⊡ Fig. 10.2.10 Beach of large pebbles on the coast at Syrianokhori, WNW of the town of Morfou.

⊡⊡ Fig. 10.2.11 Coastal slopes in weathered diabase near Pomos with a cove beach.

References Ortelius (1573) Cyprus Insular Pantazis TM (1965) Notes on the geology of southern Cyprus. Annual Report of the Geological Survey Department of Cyprus, 25 p

Thrower N (1960) Cyprus – a landform study. Ann Ass Am Geog­raph, 50 Map supplement Wilson RAM (1956) Geological Sketch Map of the Xeros Troodos area. Geological Survey Department, Nicosia

847

10.3  Syria

1.  Introduction The coastline of Syria is about 170 km long, with only a few small islands (> Fig. 10.3.1). The climate is Mediterranean, with hot dry summers and mild rainy winters. Tide ranges are small, rarely exceeding 40 cm, but there are larger fluctuations of sea level related to changes in barometric pressure, alternating onshore and offshore winds and occasional storm surges. Westerly and southwesterly winds are dominant, but the westerly fetch is limited by the proximity of Cyprus, and the southwest winds more important in producing waves and longshore drifting to the north. In summer, diurnal winds, stronger by day, produce short steep waves, and storm surges may occur in winter. In April and May, however, and also in autumn, winds are weak and the sea calm. Sea salinity is high (over 3.9%) and the water fairly warm, even in winter. The coastline is related to geological structure, the coast running parallel to the Jabal Alaouite anticline (north–south). Transverse faults, often east–west, are responsible for the main features of the steep coastline, such as the Latakie Inlet, but generally there is a narrow coastal lowland, though the shore is often rocky. Cretaceous and Tertiary limestones and marls dip seaward to outcrop on rocky coasts, but there are also unconsolidated rocks (Ras el Bassit) and volcanic formations (Banyas region). The coast is also frequently cut into Pleistocene dune calcarenite (here called ramleh), which has been used since ancient times for building stone, so that old coastal quarries are frequent. In the carbonate rocks (limestones and sandstones cemented by calcium carbonate), the generally low cliffs are fronted by shore platforms a little below high tide level, irregular in outline with almost flat surfaces. Built up on their seaward margins by algae and other organisms, these platforms, termed trottoirs, are typical of this region, and occur on about two-thirds of the coastline (> Fig. 10.3.2). Behind them, for a distance of some tens of metres, the rock is bare of vegetation and marked by deep lapies formed by solution processes.

Depositional features occur only intermittently, restricted to small bays or inlets. Sometimes, there are pebbly shores at the mouths of rivers or wadis, but more often quartzose or biogenic sands. There are often slabs of beach rock outcropping at the lower edge of the beach. Almost all the coastline is receding, even in embayments, because of the meagre supply of sand or gravel (Dalongeville 1993). Despite the erosional vigour of wadis descending from the mountains, there are only minor ⊡⊡Fig. 10.3.1 The coast of Syria.

Edited version of a chapter by Paul Sanlaville in The World’s Coastline (1985: 501–504). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

850

10.3

Syria

⊡⊡Fig. 10.3.2 Karstic dissection of a shore platform cut in Pleistocene dune calcarenite on the Syrian coast.

­ eltas, weakly convex in outline. Dunes are minor and d generally fixed by vegetation, as on the Akkar coast. Beaches and emerged platforms indicate recent oscillations of sea level of more than 2 m in the Holocene, up to the historical period. There are also examples of little deformed, but strongly uplifted, Pleistocene coastlines up to 300 m above present sea level.

2.  The Syrian coastline Cut into ophiolites, the Bassit coast is wild and picturesque. High rainfall allows oak and pine forests to extend down to the shore. The Jabal Akra majestically dominates the bay of Bassit with its sandy shores, while an outcrop of Miocene limestone forms the narrow protruding peninsula of Bassit, which is followed southward by a cliffed coast in weak rocks. Beyond that, Mesozoic limestones and Quaternary calcarenites form a series of capes and inlets bordered by corrosion platforms. The little bay of Ramlet el Beida and the Ras Ibn Hani peninsula have sheltered ancient ports, as at Ugarit, and the town of Latakia is built on a peninsula that shows remains of emerged Pleistocene coastlines.

South of Latakia, the great plain of Jableh begins, limited northward by an important structural feature, the contact of the Arabian and Anatolian plates, which has determined the northern valley of Nahr el Kebir. Except at its sandy extremities, the plain of Jableh has a rocky coast, with shore platforms developed on Pleistocene dune calcarenites of penultimate interglacial age. Between Banyas and Tartous, the coastline becomes more irregular, with some cliffs cut in Pliocene volcanic rocks. The tiny island of Rouad, the only inhabited island on the Levant coast, consists of Pleistocene dune calcarenite. The Syrian coast comes to an end in the bay of Akkar, which corresponds with a trough separating the Jabal Alaouite from the mountains of Lebanon. Developed in soft rocks (dune sands and fluvial fans, often of historical origin), the bay coastline has a long, gently curved sandy beach, the rapid recession of which is indicated by active cliffs up to 4 m high and outcrops of beach rock.

Reference Dalongeville R (1993) Recent variations on the Syrian coastline (in French). Quaternaire 4:45–53

10.4  Lebanon

1.  Introduction The coastline of Lebanon is about 230 km long (>Fig.  10.4.1). The climate is Mediterranean, with long, hot, dry summers and brief, mild, rainy winters. Tide ranges are small, up to 40 cm, but there are larger fluctuations of sea level related to changes in barometric pressure, alternating onshore and offshore winds and occasional storm surges. Westerly and southwesterly winds are dominant, with the southwest winds producing waves and longshore drifting to the north. In summer, diurnal winds that are stronger during the day produce short steep waves, and in winter, storm surges may occur. In April and May, however, and also in autumn, winds are weak and the sea is calm. Salinity is high (over 39 parts per thousand) and the water is fairly warm, even in winter (Sanlaville 1977). The coastline is related to structure: the coast runs parallel to the Lebanese Mountains (NNE–SSW), with Cretaceous and Tertiary limestones and marls dipping seaward on the western slopes, which fall more or less steeply to the coast. Transverse faults, often east–west, are responsible for the main features of the coastline, notably the Beirut peninsula. In spite of the proximity of the mountains (the culminating peak, Qornet es Saouda, 3,083 m, is only 30 km inland), the coast is generally low, though often rocky. Pleistocene dune calcarenite (known as ramleh) outcrops in low cliffs and shore platforms along parts of the coast, and has been used much since ancient times for building stone that has formed old coastal quarries. The shore platforms are submerged during high tide, but exposed during low tide as almost flat surfaces, with ramps built up on their seaward margins by algae and other organisms. Narrow benches below solution notches platforms, known as trottoirs, occur on about two thirds of the coastline. Outcrops of coastal limestone are bare of vegetation and marked by deep lapies formed by solution processes (Sanlaville 1977). Beaches of quartzose and carbonate sand are found in bays and inlets, and there is shingle at the mouths of rivers or wadis. Beach rock, formed by carbonate cementation of

beach sand, is exposed locally, often outcropping at the lower edge of the beach (Emery and George 1963). Almost all the coastline is receding, even in embayments, because of the meagre supply of sand or gravel (Dalongeville 1993). Despite the erosional vigour of wadis descending from the mountains, there are only minor ⊡⊡ Fig. 10.4.1 The coast of Lebanon.

Edited version of a chapter by Paul Sanlaville in The World’s Coastline (1985: 501–504). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

10.4

852

  Lebanon

⊡⊡ Fig. 10.4.2 Bay of Jounie, north of Beirut, Lebanon, where a steep mountain front descends to a cliffed coastline.

deltas, weakly convex in outline. Dunes are minor and generally fixed by vegetation. They occur in the Akkar and in the small bays between Beirut and Saida, as well as near Tyre.

2.  The Lebanese Coastline South of the Ne Kebit River, on the Syrian border, is a long curving sandy beach backed by dunes and cliffs in unconsolidated rock. Rapid recession of this is indicated by active cliffs up to 4 m high and outcrops of beach rock. The Nahr el Bared and the NahrAbou Ali rivers, draining from the highest parts of the Lebanese mountains, have each built convex gravelly deltas, which the sea is now trimming back. The delta of the Nahr Abou Ali has been partly derived from a sandy tombolo that links the islands of Tyrrhenian marine sandstone to the mainland in a large peninsula occupying the site of the town of Tripoli and its satellite of El Mina. Several kilometres offshore, other rocky islets appear, almost denuded of vegetation by the strong winds. In the Enfe region Cenomanian limestones form an irregular coastline bordered by a shore platform, then open to the plain of Chekka, carved in Palaeocene marls. Pebbly beaches alternate with shore platforms and rocky coasts cut in Pleistocene and dune calcarenite. Ras Chekka has the highest cliffs of the Lebanese coast, attaining 160 m and cut in Miocene limestones that lie discordantly on Palaeocene marls. South to Bahr Awali the

mountains descend steeply to cliffs up to 12 m high, and the shore is rocky and eroding. There are inlets backed by pebble beaches. Behind the Bay of Jounie the mountains are close to the coast, and a steep slope descends to cliffs and gravel beaches (>Fig. 10.4.2). The Beirut peninsula consists of marly Cretaceous limestones. It is bordered on the north and south by long sandy beaches, the southern beach being backed by large dunes on which the airport has been built. Further, south limestone headlands separate a series of sandy beaches at Damour, Jiyeh and Rmeileh. Beyond the town of Saida the coastal plain reappears, widening southward, and sandy beaches are often accompanied by extensive outcrops of beach rock. The sandy tombolo of Tyre, formed within the historical period, is surmounted by dunes that have been much disturbed by man, and have overrun the site of the Roman hippodrome. To the south the coast is again cliffed and rocky, with prominent notches and trottoirs on the limestone cliffs toward Rosh Haniqra on the Israel border.

References Dalongeville R (1993) Recent variations on the Syrian coastline. (in French) Quaternaire 4:45–53 Emery KO, George CJ (1963) The shores of Lebanon. American University of Beirut, Miscellaneous Papers in Natural Science, 1:1–10 Sanlaville P (1977) Geomorphological study of the coastal region of Lebanon. Université Libanese, Section des Etudes Géographique, Beirut

10.5 Israel (with the Gaza Strip)

Yaacov Nir

1.  Introduction The combined coastline of Israel and the Gaza Strip is 230 km long (>Fig.  10.5.1). It is of generally low relief, except for mountains of Cretaceous (Cenomanian) rock that reach the coast at Rosh Haniqra and Mount Carmel. It has beaches that are partly supplied with sand from the River Nile and erosion of its delta and offshore shoals. This has been carried northward by longshore drifting, and augmented by calcareous sand from the sea floor and eroding cliffs. The beaches are backed by dunes and dune calcarenite, carbonate cemented quartz sandstone, known locally as Kurkar (>Fig. 10.5.2), which in places has been trimmed back in cliffs and shore platforms. Harder calcarenite layers formed by carbonate precipitation on or beneath former land surfaces may protrude from cliffs (>Fig. 10.5.3). Dune calcarenite parallel ridges of Pleistocene age run along much of the coastline (apart from Haifa Bay), extending up to 20 km inland in southern Israel, decreasing in island distance northward to a few hundred metres along mt. Carmel Coastal plain. The coastal climate is Mediterranean, with typically hot, humid summers and mild, rainy winters. Summer winds are gentle, but winter storms are occasionally severe. Winds are mainly southwesterly or northwesterly. Rainfall is up to 350 mm/year in the south, 500 mm/year on the central coast, and 600 mm/year north of Haifa. Tides are semi-diurnal, with a range of 30–60 cm. There are seasonal fluctuations with a range of 20 cm, the  highest levels occurring at the end of summer and the  lowest at the end of winter. The wave climate has three  wave seasons: the winter months (December– March) have the highest mean significant wave heights ­(1.2–1.6  m), the  lowest monthly mean significant wave heights occur in May and from October to early November, and the summer months (June through September) are intermediate, with 0.5–1.0 m waves. The longest fetch is for westerly waves from the Straits of Sicily (about 2,400 km), but the island of Cyprus forms a slight wave shadow to the northern israel Coast, effective during spring and autumn when west northwesterly

winds are dominant. In general, high waves are from west southwest through west to north northwest, generated by depressions passing through the Mediter­ranean from west to east. Beach rock occurs extensively on the Israeli coast, with the exception of Haifa Bay (Emery and Neev 1960). It has a gentle seaward inclination and strikes parallel to the coastline (>Fig. 10.5.4). It is exposed where the overlying beach sand has been removed by erosion. Pumice occurs as grey fragments that are usually well rounded and up to 15 cm in diameter. They are very light and float in the sea, and have probably come from the Greek volcanoes, notably Santorini. ⊡⊡ Fig. 10.5.1 The coast of Israel and the Gaza Strip.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_10.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

854

10.5    

Israel (with the Gaza Strip)

⊡⊡ Fig. 10.5.2 Cross-bedded dune calcarenite, of the Kurkar coastal cliff at the Herzliya coast.

⊡⊡ Fig. 10.5.3 A protruding layer of hard calcarenite where softer underlying soft red loam (locally know as “hamra”) has been removed.

2.  The Coastline From Rosh Haniqra southward for 20 km, much of the coastline is backed by Pleistocene dune calcarenite, bordered by shore platforms or beaches of calcareous sand. Beach rock is extensive. Six short rivers draining the Western Galilee Mountains deliver water and sediment to this coast, and are the source of mostly limestone pebbles on beaches and cemented in beach rock (beach conglomerate).

To the south is the Acre promontory, which is the site of an ancient harbour known as Ptolemais. The fine sandy beaches contain relatively high proportions of carbonate, derived from the reefs of submerged dune calcarenite on the sea floor. Acre marks the limit of quartz sand that has drifted northward from the Nile (Masters 1996). Haifa Bay to the south is backed by a coastal plain up to 10 km wide, and has a crescentic dune field extending up  to 2 km inland. It occupies the seaward end of the

Israel (with the Gaza Strip)  

   10.5

⊡⊡ Fig. 10.5.4 Layers of beach rock dipping seaward on the shore south of Hadera Beach. 

⊡⊡ Fig. 10.5.5 Sand accumulation beside the rim breakwater at Bat-Yam, south of Tel Aviv indicates local alternations of longshore drifting.

Zevulun Valley graben. Beaches widen towards the centre of the bay as the result of convergent longshore drifting. The coast curves out to the Haifa promontory, at the northern end of the Mount Carmel Range. Haifa Bay has two major rivers, the Qishon and Na’aman, which deliver mainly fine sediment to the shore. Dune calcarenite and beach rock are absent from this part of the coast. Haifa Harbour was built in 1929–1932 at the northern end of the Nile derived longshore drift, Nile derived longshore drift

ends at Haifa Bay that plays as the final sink for the imported sediment. The main breakwater of the Haifa Harbour accumulated since its construction an enormous amount of sand, pre­venting the supply to Haifa Bay beaches. From Cape Carmel, a narrow coastal plain widens southward as the mountain range recedes. Its seaward fringe consists of one to three ridges of dune calcarenite up to 15 m high, running parallel to the coastline and interrupted by artificially cut river courses. The ridges are separated by swale plains with dark terrestrial clay soils.

855

856

10.5    

Israel (with the Gaza Strip)

South to the ‘Atlit promontory the beaches have scanty deposits of sand, which generally forms a thin veneer over a rocky shore. Limestone pebbles brought down by short rivers from the Carmel Mountains are mainly exposed during winter.

The fact that Nile-derived beach sand extends north to Haifa indicates long-term northward drifting, but there are occasional episodes of southward longshore drifting between Haifa and Tel Aviv (Golik 1993; Klein et al. 2007; Soshani et al. 1996). These have resulted in ⊡⊡ Fig. 10.5.6 A peninsula at Nakhsholim, north of Caesarea, where a small island has been attached to the coast by a sandy tombolo.

⊡⊡ Fig. 10.5.7 The modern harbour at Caesarea is built over the site of King Herod’s ancient harbour.

Israel (with the Gaza Strip)  

⊡⊡ Fig. 10.5.8 Dune calcarenite cliffs south of Netanya, with a narrow sandy beach and huge Kurkar blocks at the base of the ca. 40 m high cliff.

   10.5

attacked by atmospheric weathering and erosion at their base during winter storms, and are subject to landslides (>Fig. 10.5.9). Incised V-shaped valleys have been truncated by cliff recession, and in some cases end as hanging valleys above beach level. The beaches are composed of medium to fine quartz sand with calcareous sand and pebbles from the dune calcarenite and pelecypod shells of local origin. At Netanya the sandy beach is periodically eroded by storms, exposing underlying gravel and rock outcrops. The beaches between Tel Aviv and Gaza are sandy, mostly wide, backed by sand dunes and patchy dune calcarenite. There are numerous artificial structures such as harbours, marinas, power-station cooling basins, detached breakwaters, groynes, and sea walls (Nir 1988). These have led to erosion of beaches and cliffs in sectors where the longshore drift has been cut off (>Fig. 10.5.10).

3.  Gaza Strip

sand accretion on both sides of some breakwaters (>Fig.  10.5.5). The 38 km coastal sector between Cape Carmel and Caesarea has narrow beaches, mainly in coves in the dune calcarenite fringe. A few smaller dune calcarenite relics form small islets and reefs close to the coast. Some have been attached to the mainland by sand deposition (>Fig. 10.5.6). Locally on the beach there are outcrops of black swampy clay, deposited on the present coastal plain and the emerged sea floor during low sea level (glacial) phases of the Pleistocene. An ancient harbour lies beneath the modern port (>Fig. 10.5.7). The coast from Caesarea to the northern end of the Gaza strip is 108 km long. From Caesarea south to Tel Aviv, the Sharon Plain ends in high dune calcarenite cliffs behind beaches that are narrow, and often ­disappear during winter storms (>Fig.  10.5.8). These cliffs are

The 40 km long Gaza Strip coastline is smoothly curved in  the southeastern section of the Levantine Basin. The northern section is backed by cliffs of dune calcarenite rising at least 20 m above the shore, with beaches that widen southward. The beaches have been supplied by strong northward drifting of sand from the Nile delta, and are backed by generally mobile dunes. These beaches have been depleted by sand mining, an activity that has been practiced for many years. It has resulted in accelerated erosion of backshore cliffs. A provisional harbour that was constructed in Gaza in the early 1970s consisted of two 120 m long groynes 500 m apart, projecting at right angles to the coastline, which resulted in accretion on the southern side and severe erosion of beaches and coastal land to the north. Short sea walls were constructed to protect these beaches, but they merely transferred the erosion northward. South of Gaza the cliffs cut in dune calcarenite are lower, and the lightly cemented Pleistocene quartzose sandstones are  soft enough in places to form sloping bluffs rather than cliffs. Offshore, a veneer of quartz sand up to 5 m thick covers a rocky surface of planed-off dune calcarenite or beach rock, layers of which outcrop on the sea floor to a depth of 4.5 m. There has been accretion on the southern beaches of the Gaza Strip. Towards the Egyptian border at Rafah the beaches are 40–50 m wide, transitional to the very wide Sinai beaches, which reach widths of 100 m.

857

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10.5    

Israel (with the Gaza Strip)

⊡⊡ Fig. 10.5.9 Landslides on dune calcarenite cliffs form debris cones that are removed by wave action.

⊡⊡ Fig. 10.5.10 Erosion at Ashqelon, downdrift of detached breakwaters that have trapped northward­drifting sand, resulting in a beach deficit.

References Emery KO, Neev D (1960) Mediterranean beaches of Israel. B Geol Surv Israel 26:1–24 Golik A (1993) Indirect evidence for sediment transport on the continental shelf of Israel. Geo-Mar Lett 13:159–164 Klein M, Zvieli D, Kit E, Shteinman B (2007) Sediment transport along the coast of Israel: examination of fluorescent sand tracers. J Coastal Res 23:1462–1470

Masters PM (1996) Paleocoastlines, ancient harbours, and marine archeology. Shore Beach 64:8–17 Nir Y (1988) Israel. In: Walker HJ (ed) Artificial structures and shorelines. Kluwer Academic Publishers, Los Angeles, CA, pp 253–260 Shoshani M, Golik A, Degani A, Lavee H, Gevirtzman G (1996) New evidence for sand transportation along the coastline of Israel. J Coastal Res 12:311–325

11.0 C  aspian and Aral Seas – Editorial Introduction

1.  Introduction The Caspian and Aral Seas occupy basins in the west of Central Asia. Both have shown major oscillations of sea level, independent of those in the world’s oceans, and providing opportunities to study the response of coastal processes and landforms to rising and falling sea levels (Dolotov and Kaplin 2005). The Caspian Sea level fell between 1929 and 1977, and has since risen again, while the Aral Sea has been drastically reduced since 1961.

2.  Caspian Sea Coast The Caspian Sea (> Iran Caspian, > Turkmenistan, > Kazakhstan, > Russian Caspian, > Azerbaijan) was formerly linked to the Atlantic Ocean by way of the Sea of Azov, Black Sea, and Mediterranean, but the link was broken in Pliocene times. Since then, the sea has occupied a composite basin bordered by the Quaternary Russian Plain to the north and mountain ranges to the south. The northern lowlands include the large Volga delta and the swampy plains to the east. The Caucasus Mountains, consisting of Mesozoic and Tertiary formations folded by Alpine earth movements, extend south-eastward toward the coast between Makhachkala and Baku, and are bordered by a narrow coastal lowland with minor cliffs. The Kura delta occupies the Kura-Araks depression south of the mountains, and beyond it another range, the anticlinorial Talysh Mountains, runs south-east toward the coast between Lenkoran and Rasht. Again, the mountains are bordered by a narrow coastal lowland. The southern (Iranian) coast of the Caspian Sea (> Iran Caspian) is bordered by a lowland up to 30 km wide, backed by the Elburz Mountains. The lowland widens in Turkmenistan on the south-east coast of the Caspian Sea, where Tertiary formations outcrop in a low plateau extending north to the Kora-Bogaz-Gol, an almost enclosed bay. To the north, the Tertiary formations rise to

the Mangyshlak hills, which run NW to the Shevchenko Peninsula. The northern part of the Caspian Sea is shallow, and there were major changes in the extent of the land as a result of sea level lowering and revival. The Caspian Sea tides are small (a few cm), but there are short-term variations in level of up to 3 m as the result of wind action, changes in barometric pressure, and the formation of seiches (periodic oscillations of water level in response to sudden changes of pressure). Its salinity is generally low, averaging about 1.3%, but increases to more than 30% in the Kora-Bogaz-Gol because of high evaporation. Wave action on the west coast is predominantly from the SE, on the south coast from the N, and on the east coast from the SW.

3.  Aral Sea Coast The >Aral Sea occupies an oval-shaped tectonic depression about 500 km to the east of the Caspian. It formed originally in the Cretaceous, and has been shaped by Tertiary and Quaternary sedimentation. Uplifted Cre­ taceous sandstones and limestones outcrop in hilly areas to the west, and in spurs forming promontories on the north coast. The low-lying east and south coasts are dominated by the large deltas of the Syr Darya and Amu Darya Rivers. Lowering of sea level since 1961 has exposed wide emerged plains with stranded beach ridges. Wave action is related to the prevailing winds from the NE in autumn and winter, and from the W and SW in spring and summer. Salinity is generally low, averaging 1.0% but increasing in the SE, particularly in sheltered bays.

Reference Dolotov Y, Kaplin P (2005) Coastal zone of the Caspian Sea. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 198–203

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

11.1 Iran – Caspian Sea Coast

1.  Introduction The Caspian coast of Iran is about 740 km long, stretching like a horseshoe around the southern side of the Caspian Sea. The land descends from the lower slopes of the Elburz Mountains to the Caspian Sea, now about 26 m below global mean sea level. The surface of the Caspian Sea has shown long-term oscillations, and fell about 3 m between 1930 and 1977, since when it has risen about 2.5 m, providing an important example of a modern marine transgression (>Fig. 11.1.1). The coastal plain varies in width from less than 2 km in parts of the western and central sections to more than 40 km in the delta regions and in the Turcoman Steppe to the east. Geologically, it is of Late Pleistocene to Holocene age. The Elburz Mountains are fronted by steep aprons of coarse gravel and sand that decline to an alluvial coastal

plain. Discharge of water and sediment from the many rivers (>Fig.  11.1.2) that flow down to the Caspian has produced a swampy lowland, parts of which have been converted into ricefields. Construction of dams on rivers has reduced fluvial sediment yields (Voropaev et al. 1998). The climate is strongly seasonal, with hot wet summers and cool to cold damp winters. Ramsar has 34°C in July and 2°C in January, and rainfall diminishes from 1,834 mm at Bandar Pahlavi in the west to 1,265 mm at Ramsar and 807 mm in Babolsar to the east. The Elburz foothills are extensively forested, and the coastal plain is well vegetated except where dunes are bare and mobile. Northerly winds prevail, and recurrent NW storms generate a dominance of eastward longshore drifting along the southern coast. Tide ranges are very small, but sea level can rise and fall by up to 2 m as the result of storm surge events, ­fluctuations

⊡⊡ Fig. 11.1.1 Changes in the level of the Caspian Sea from 1930, as indicated on the Baku tide gauge (Azerbaijan). Sea level fell intermittently until 1977, then rose fairly rapidly, probably because of a trend towards a more humid climate and increased river runoff in the region, but possibly related to changes that have occurred in the Kara Bogaz Gol (Turkmenistan). (Courtesy S. Lukyanova.)

Edited version of a chapter by E. Ehlers in The World’s Coastline (1985: 487–490). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

11.1

862

Iran – Caspian Sea Coast

⊡⊡ Fig. 11.1.2 The River Shahi is one of several rivers draining from the Elburz Mountains that deliver sand and gravel to the Caspian coast. (Courtesy Geostudies.)

⊡⊡ Fig. 11.1.3 Changing levels of the Caspian Sea, A – during Late Pleistocene times, B – between 500 bc and 1977 ad.

A 45 - 50 m.s.l. 35 m 25 m

14 m 0m –10/–12 m below m.s.l. –16/–18 m

Present level of the Caspion Sea Middle Chvalynian Stage

Early Chvalynian Stage

Late Chvalynian Stage Holocene

Neo-Caspian Stage

B 22 m below m.s.l.

26 m Present level of the Caspian Sea (–28 m below m.s.l.)

– 30 m below m.s.l. 500 BC

0

500

1000

1500

2000 AD

Iran – Caspian Sea Coast

11.1

⊡⊡ Fig. 11.1.4 Pleistocene terraces indicating former higher levels of the Caspian Sea.

⊡⊡ Fig. 11.1.5 The barrier on the south coast of the Caspian Sea showing levels 2 m, 4 m and 6 m above the 1977 level.

in barometric pressure, heavy rainfall and river inflow, and major oscillations (seiches) in the Caspian Sea. Salinity varies around an average of about 12 parts per thousand, and water temperatures range from about 25°C in summer down to 8–11°C in winter. Oscillations of the Caspian Sea during Late Pleistocene and Holocene times have been related to climatic fluctuations (>Fig. 11.1.3). On the slopes fronting the Elburz Mountains there are terraces (>Fig.  11.1.4) that indicate sea levels higher than the present in the Late Pleistocene (Würm or Wisconsin glacial stages). The terraces, which extend around the Caspian coast, can be correlated with fluvial terraces in valleys descending the north slopes of the Elburz Mountains, which in turn can be related to moraines of the last glaciation. There is evidence that the Caspian fell about 50 m below global mean sea level in the early Holocene Mangyshlak stage, about 3,500 years ago, before the Neo-Caspian marine transgression. This rose above present level, and

formed a barrier with beaches 2, 4 and 6 m above the 1977 level of the Caspian Sea (>Fig. 11.1.5). While the highest level is partly covered by recently formed dunes, the lower levels are separated from each other by swampy depressions, partly occupied by deflected river mouths (Fedorov and Skiba 1960). The lowering of sea level between 1930 and 1977 re­­ sulted in the formation of wide sandy beaches (>Fig. 11.1.6) and derived dunes (>Fig.  11.1.7), and the incision and seaward and extension of rivers. The subsequent sea level rise has resulted in submergence of beaches and cliffing of dunes, as well as the revival of lagoons in the marshy area landward of the transgressive Holocene coastal barrier. The delta of the Safid River, forming the Gilan coastal plain, is cut through old beach ridges formed when the Caspian stood at higher levels, and includes lagoons, shallow bays and marshes (>Fig.  11.1.8). Krasnozhon et al. 1999 have traced delta evolution using space imagery. The Haraz, Babol, Talar, and Tejan rivers further to the

863

11.1

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Iran – Caspian Sea Coast

⊡⊡ Fig. 11.1.6 The broad emerged sandy beach at Babolsar, colonised by shrubby vegetation. (Courtesy Geostudies.)

⊡⊡ Fig. 11.1.7 Dunes spilling inland behind the emerged beach on the south coast of the Caspian Sea west of Babolsar. (Courtesy Geostudies.)

east, form the Mazanderan plain, with channels bordered by levees. River flooding occurs frequently on the coastal plain, especially during winter rains and the melting of snow in the Elburz Mountains in spring (Raynal 1979). The Holocene coastal barrier is backed by marshy ­lowlands, most of which are situated below mean sea level, formed mainly by silt and clay with a considerable content of salt. The water table is at shallow depth, ­typically

about 0. 6 m, and groundwater is fed both by lateral seepage from the rivers and by aquifers in the gravelly fans fronting the Elburz foothills.

2.  The Caspian Coast of Iran South from the Azerbaijan border a narrow coastal plain fronts a mountain range. The coastal plain widens

Iran – Caspian Sea Coast

11.1

⊡⊡ Fig. 11.1.8 The river mouth at Babolsar, partly blocked by sand spits in the dry season. (Courtesy Geostudies.)

⊡⊡ Fig. 11.1.9 Emerged beach ridge in front of the escarpment that borders the Elburz Mountains east of Sari. (Courtesy Geostudies.)

s­ outhward to Karganrud, and is broad in Gilan province as the coastline swings east. Bandar Anzali stands beside a gap in the barrier spit that encloses a large lagoon. To the east is the Safid delta with several distributaries, ending in a lobe near Dehgan, where the coast turns south, then ESE.

The coastal plain narrows in front of a steep mountain slope past Ramsar, and narrows again at the mouth of the Shahsavar River valley (>Fig.  11.1.9). The Bud-e-Chalu River flows from a similar valley across the plain to Now Shaht, and the coast swings ENE at Suledeh. The coastal plain widens behind Mazandaran and Mahmudabad as

865

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Iran – Caspian Sea Coast

the mountain front recedes, and at Babolsar the river mouth is often constricted by the growth of a spit from the western side (>Fig. 11.1.8). The coastal lowland is backed by emerged coastline features (>Fig. 11.1.9). The lowland behind the coastal barrier becomes a lagoon, Gorgan Bay, behind a spit that shows evidence of recurves indicating intermittent eastward growth. This has deflected the lagoon outlet close to Bandar-e Shah, and the coastline curves northward in front of the wide dry plains of Terabad. Alexander’s Wall stops short of the present coastline, and to the north the Solanchak River flows intermittently down to the eastern coast. The desiccated marshy plains extend across the border into Uzbekistan.

References Fedorov PV, Skiba LA (1960) The sea-level fluctuations of the Caspian and the Black Sea during the Holocene (in Russian). Izv Akad Nauk SSSR Ser Geogr 4:24–34 Krasnozhon GF, Lahijani H, Voropayev GV (1999) Evolution of the delta of the Sefidrud River, Iranian Caspian Sea coast, from space imagery. Mapp Sci Rem Sens 36:256–264 Raynal R (1979) An example of accelerated erosion in a temperate region without a dry season: Lahidjan, coastal zone of the Caspian Sea. Comptes-rendus: colloque, érosion agricole des sols en milieu tempéré non méditerannéen, Strasbourg-Colmar, pp 39–42 Voropaev GF, Krasnozhen GF, Lachojani HK (1998) River drainage and stability of the Iranian coast of the Caspian Sea (in Russian). Water Resour 25:747–758

11.2 Turkmenistan

Andrei Selivanov*

Introduction The coastline of Turkmenistan (>Fig. 11.2.1) is 1,768 km long. The Republic of Turkmenistan has extensive subtropical desert areas, and a climate of hot summers and cold winters. The coast is generally depositional, with formations that accumulated in Quaternary times, and, in the absence of rivers to supply fluvial sediment, the beaches are dominated by shelly material and oolites swept in from the sea floor (Leontiev 1985; Svitoch and Yanina 1997), with local accessions from cliffs. The morphology of the Caspian Sea is related to its geology, with the Derbent depression (780 m deep) and the South Caspian depression (1,022 m deep) corresponding with structural troughs. There are fluctuations of up to 3 m in sea level, related to wind action and barometric pressure. Tide ranges are small, only a few centimetres. On the east coast of the Caspian Sea, westerly winds generate the dominant waves, which are up to 6 m high during storms. Waves that arrive from the southwest cause northward longshore drifting, while waves from the northwest cause southward longshore drifting, dominant in Turkmenistan. Separated from the oceans in Lower Pliocene times (Svitoch et al. 1998), the Caspian Sea has shown fluctuations related to climatic changes within the bordering drainage basins. During the Quaternary there were re­­peated glaciations of the Russian Plain and the Caucasus Mountains, and intervening deglaciations. The level of the Caspian Sea consequently rose and fell in the Pleistocene, attained a level 50 m higher than at present in the Middle Pleistocene, and falling 50 m below present level in the Late Pleistocene. Marine terraces correspond with stillstands during these fluctuations. On the relatively stable eastern coast, there are terraces up to 70 m above present level (Rychagov 1997). During the Holocene sea level rising, the ­Neo-Caspian marine transgression, culminated in the eighteenth century, when it reached 6 m above the present sea level, about 26 m below global mean sea level. Since then there has been a substantial lowering, with a sharp fall of about 3 m between 1930 and 1976, which reduced the area of the Caspian Sea by more than 40,000 sq. km. During the sea

level fall, there was emergence of land margins and shoals became islands, along with shoreward drifting of sand from the sea floor on to beaches, spits, and barriers. In 1977, when the sea level had declined to its lowest position (−29.02 m), the level lowering came to an end, and was followed by a modern marine transgression of about ⊡⊡ Fig. 11.2.1 The coastline of Turkmenistan. (Courtesy Geostudies.)

*Edited version of chapter 11.2 (Turkmenistan) in The World’s Coasts: Online (2003) by Andrei O. Selivanov. All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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11. 2    

Turkmenistan

2.5 m by 1995, probably because of a trend towards a more humid climate and increased river runoff in the region, but possibly related to changes that have occurred in the Kara Bogaz Gol (>Fig. 11.2.1). This has caused a general retreat of the coastline, flooding and waterlogging of agricultural land, intensification of storm surge effects, and revival of wave erosion on previously abandoned cliffs. The proportion of eroding coastline has increased from 7% in 1978 to 22%. The low gradients of the coastal land resulted in marine submergence generating narrow (about ­200–300 m) and extensive (several kilometres in length) depositional barriers fronting shallow lagoons, all transgressive as sea level rise proceeded. The implications of these changes for other coasts in the expected global sea level rise have been extensively discussed (Ignatov et al. 1993; Selivanov 1997; Kaplin and Selivanov 1999; Rychagov 2002). The southern (Iranian) coastline of the Caspian Sea extends to Bandar-e Shah, then turns abruptly northwards, past the mouth of the Gorgon River to the Turk­ menian border at the mouth of the Atrek River. The smooth coast is low and sandy, with beaches backed by old dunes and a vast saline (solonchak) plain where former lagoons have dried out. As sea level has fluctuated, some of the lagoons remain as dry solonchaks, while others have revived as saline lagoons. Erosion has cut back into Pleistocene deltaic clays around the mouth of the Atrek River. Numerous and very large active mud volcanoes are the most picturesque features in the coastal land of this region. Some of them are located near the shoreline. The hinterland remains dry and sandy to the north, as far as Turkmenskiy Bay. This marks the beginning of the curious irregular protrusion of the Cheleken Peninsula, a ridge of Neogene limestone and sandstone, which ends in a rapidly eroding west-facing cliffed coast (>Fig. 11.2.2). Two large sand spits have grown southward and northward from these cliffs, forming a so-called “wing cape”. This is a coast that sometimes receives southwesterly wave action, causing northward drifting, and, at other times, northwesterly wave action, causing southward drifting. These spits have been influenced by changes in sediment supply that accompanied variations in sea level regime. The big aeolian features cover not only the sediment spits, but also almost all the surface of the peninsula. The peninsula became an island during the sea level rise in 1978–1995, but then rejoined the land during a slight emergence. To the south is the long, narrow island of Ogurchinskiy, a shelly barrier beach growing southward as the result of longshore drifting, which becomes dominant as the southwesterly fetch diminishes.

⊡⊡ Fig. 11.2.2 Active cliff and narrow sand beach at the westernmost part of the Cheleken Peninsula. (Courtesy R. Kurbanov.)

North of the Cheleken Peninsula is the large, irregular Krasnovodsky Bay, its mouth partly blocked by spits, and the island of Kizyl-Su, dominated by sandy sediment ­carried in from the sea floor. In recent decades the spit has been retreating at more than 3 m/year as the result of wave erosion. The bay has shallow offshore areas and extensive muddy flats in its inner parts. It is bordered by a  wide zone that is subject to flooding during westerly storm surges. The port of Turkmenbashi City (formerly Krasnovodsk) stands on the northern shore of this bay, backed by hilly topography which extends to the coast near Kuuli-Mayak. The Krasnovodsk Plateau to the north is composed of Pleistocene shelly limestone, which forms several small coastal promontories subject to intensive erosion during periods of rising sea level. Some sectors advanced at a rate of 30–35 m/year during the sea level fall of 1930–1978, and have retreated at a much lower rate during the subsequent sea level rise. To the north is the almost landlocked Kara Bogaz Gol, a large embayment (18,000 sq km), which had only a

Turkmenistan  

­ arrow outlet through an enclosing barrier (7–10 m high) n of shelly sand formed from sediment swept in from the sea floor. Kara Bogaz Gol means Black Throat Bay in the Turkiс language. It is a hypersaline lagoon bordered by salt flats and tall cliffs. Large quantities of water (up to 8–10 km3/ year) used to flow in from Caspian Sea and evaporate from its surface, causing salinity to rise to 300 parts per thousand and contributing to the lowering of sea level in the Caspian. A waterfall 3 m high existed in the outlet channel over the limestone ridge. In order to counter losses of water from the Caspian Sea, a 7 m high dam was built to close the entrance. Evaporation losses duly diminished, but the hypersaline lagoon continued to dry out and extensive solonchaks formed on the bottom of the Kara Bogaz Gol. The salt layer is up to 3 m thick, and is a source of several unique chemical compounds, including the mineral mirabilite, which is quarried. The sea level began to rise in the late 1970s, and in 1992 the rulers of the newly independent Turkmenistan destroyed the dam, allowing Caspian water to resume flow into the Kara Bogaz Gol. Evaporation of inflowing water may have contributed to the slight fall in level of the Caspian Sea from 1995 onwards. Turkmenistan officials estimate that evaporation  from the revived Kara Bogaz Gol lowers the sea level  by 6–8 cm/year (UNEP 1997), but this is an overestimate. On the northern side of

   11. 2

Kara Bogaz Gol, hilly country runs out to Cape Suz, which marks the border with Kazakhstan.

References Ignatov YI, Kaplin PA, Lukyanova SA, Solovieva GD (1993) Evolution of the Caspian Sea coasts under conditions of sea level rise: model for coastal change under increasing “Greenhouse Effect”. J Coastal Res 9:50–57 Kaplin PA, Selivanov AO (1999) Sea-level changes in Russia and coastal  evolution: past, present and future (in Russian). GEOS, Moscow Leontiev OK (1985) Caspian USSR. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 481–486 Rychagov GI (1997) Pleistocene history of the Caspian Sea (in Russian). Moscow University Press, Moscow Rychagov GI (2002) The Caspian Sea: past, present and future (in Russian). Zhyvopisnaya Rossia 3:39–42 Selivanov AO (1997) Coastal modifications of the Caspian Sea and other Central Asian lakes as natural models for coastal responses to the global sea-level rise. Bollettino di Geofisica 37(1):114–129 UNEP (United Nations Environmental Programme) (1997) Conse­ quences  of climate changes in the Caspian Sea in the Caspian Sea region: regional review document. UNEP, Geneva Svitovch AA, Yanina TA (1997) Quaternary sediments of the Caspian Sea (in Russian). Lomonosov Moscow State University Publication, Moscow

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11.3 Kazakhstan

Andrei Selivanov*

1.  Introduction The coastline of Kazakhstan (> Fig. 11.3.1) is 1,894 km long. The Kazakh Republic has a steppe and desert hinterland, with a continental climate: warm to hot summers and cold winters. Much of the coast is depositional, with formations that accumulated in the Quaternary, and, in the absence of rivers to supply fluvial sediment, the beaches are dominated by shelly material and oolites swept in from the sea floor (Svitoch and Yanina 1997), with local accessions from cliffs. ⊡⊡ Fig. 11.3.1 The Kazakhstan coastline. (Courtesy Geostudies.)

The morphology of the Caspian Sea is related to its geology, with the Derbent depression (780 m deep) and the South Caspian depression (1,022 m deep) corresponding with structural troughs. There are fluctuations of up to 3 m in sea level, related to wind action and barometric pressure. Tide ranges are small, only a few centimetres. On the east coast of the Caspian Sea westerly winds generate the dominant waves, which are up to 6 m high during storms. Waves that arrive from the southwest cause northward longshore drifting, while waves from the northwest cause southward longshore drifting. Separated from the oceans during the Lower Pliocene (Svitoch et al. 1998), the Caspian Sea has shown fluctuations related to climatic changes within the bordering drainage basins. During the Quaternary there were repeated glaciations of the Russian Plain and the Caucasus Mountains, and intervening deglaciations. The level of the Caspian Sea consequently rose and fell in the Pleistocene, attaining a level 50 m higher than at present in the Middle Pleistocene, and falling 50 m below present level in the Late Pleistocene. Marine terraces correspond with stillstands during these fluctuations. On the relatively stable eastern coast there are terraces up to 70 m above present level (Rychagov 1997). During the Holocene a sea level rise, the Neo-Caspian marine transgression, culminated in the eighteenth century, when it reached 6 m above its present level, about 26  m below global mean sea level. Since then there has been a substantial lowering, with a sharp fall of about 3 m between 1930 and 1976, which reduced the area of the Caspian Sea by more than 40,000 sq. km (> Fig. 11.3.2). During the sea level fall land margins emerged, shoals became islands, and there was shoreward drifting of sand from the sea floor on to beaches, spits, and barriers. In 1977, when the sea had declined to its lowest level (–29.02 m), the level lowering came to an end, and was followed by a modern marine transgression of about 2.5 m by 1995. This was probably due to a trend towards a  more humid climate and increased river runoff into the Caspian, but was also possibly related to changes that have occurred in the Kara Bogaz Gol in Turkmenistan. The ­rising sea level has caused a general retreat of the coastline, flooding and waterlogging of agricultural land, intensification of storm

*Edited version of chapter 11.3 (Kazakhstan) in The World’s Coasts: Online (2003) by Andrei O. Selivanov. All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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11.3

Kazakhstan

⊡⊡ Fig. 11.3.2 Coastline changes in the northern Caspian Sea between 1930 and 1976 resulting from a sea level fall of about 2.67 m. (Courtesy O.K. Leontyev.)

surge effects, and revival of wave erosion on previously abandoned cliffs. The proportion of eroding coas­­tline has increased from 7% in 1978 to 22% at present. The low gradients of the coastal land resulted in marine sub­­mergence generating narrow (about 200–300 m) and extensive (several kilometres in length) depositional barriers fronting shallow lagoons, all transgressive as sea levels continued to rise. The implications of these changes for other coasts in the expected global sea level rise have been extensively ­discussed (Ignatov et al. 1993; Selivanov 1997; Kaplin and Selivanov 1999; Rychagov 2002).

2.  The Kazakhstan Coastline The low hilly coastline north of the Turkmenian border continues behind Kazakhskiy Bay, which contains Ken­ derli, a barrier spit that has grown northward, indicating a prevalence of longshore drifting in that direction on the northeast Caspian coast. Shelly beaches are extensive on this coast, derived from inwashed sea floor sediment. Oolite sands and shelly sandstones are also present in a zone of old beach formations up to 2 km wide and 15 m above the present Caspian Sea. On the north coast of

Kazakhstan

the  bay is a cuspate promontory, Mys Rakushechnyy (which means Shelly Cape), after which the coast declines into Aleksandr Bay. To the west is a depositional foreland,  Peschany Point (which means Sandy Point), with dunes surmounting beach ridges (> Fig.  11.3.3). Both forelands  have triangular hinterlands of Neogene and Pleis­tocene shelly sandstone, cliffed during phases of high sea level and bordered by Holocene beach ridges of shelly sand. The coast then swings north to Aktau City (formerly Shevchenko). The Mangyshlak Plateau (up to 70 m high) on Miocene carbonate rocks runs northwest behind Aktau and out to  the broad Tyub Karagan peninsula on which Fort Shevchenko stands. This peninsula formerly had cliffs cut in Miocene limestone, clay, and marl, which were left stranded as bluffs when the sea level fell, but the rising sea has now returned to them. There are extensive landslides on cliffs cut into clays and marls, particularly on the north coast of the Tyub Karagan peninsula. During the sea level rise until 1929, these cliffs were cut back at least 2 m/year, but when the sea level fell between 1930 and 1977, this erosion stopped, and beaches up to 40 m wide formed in front of subaerially degrading cliffs. When sea level rose again after 1978, the beaches dwindled and cliff recession

⊡⊡ Fig. 11.3.3 The east Caspian coast between the capes of Rakushechnyy and Peschany. (Courtesy E.N. Badyukova.)

11.3

revived. There is severe coastal erosion around the cities of Aktau and Fort Shevchenko (> Fig. 11.3.4). To the north is an inlet, Mangyshlakskiy Bay, and then the low, undulating Buzachi Peninsula, which has extensive salt marshes. The northeastern section of the Caspian Sea is an area in which extensive changes occurred as the result of the lowering of sea level between 1930 and 1976. Shoals of shelly sand emerged to form the Tyuleni (Seal) Islands, and barrier beaches along the 1930 coastline were left stranded inland on the Buzachi Peninsula. The 1930 coastline runs behind a broad bay, the floor of which emerged as the marshy Mertvyy Kultuk and Peski Karakum lowlands (>Fig. 11.3.3). The fall in sea level exposed more than 20,000 sq. km, much of which dried out as saline flats (solonchaks). Emergence led to the formation of shalygs, offshore sand ridges of barchan-like form, shaped by the interaction of waves and currents in a shallowing sea. The Emba River and the Ural River extended their courses across the emerged sea floor, the Ural (depleted by the use of its water for irrigation) building a small delta at Gur’yev City. During the sea level fall (1930–1977), the delta front advanced by up to 40 km. Clay dunes formed on the emerged arid coastal plain, and the low sandy coastline is now fronted by wide tidal flats. During the modern marine

873

874

11.3

Kazakhstan

⊡⊡ Fig. 11.3.4 The active cliff of Miocene limestone near Aktau City. (Courtesy T. Yanina.)

transgression which began in 1978, extensive parts of these emerged lowlands, including solonchaks, marshes, shalygs, and clay dunes, as well as the margins of the Ural River delta, have passed again beneath the sea. In addition, storm surges can inundate low-lying areas up to 60 km inland. They scour the sea floor and sweep silt deposits landward. Oil and natural gas exploration has led to some pollution, and dams and other structures built to halt erosion and prevent sea flooding need maintenance and, in some cases, protective beach nourishment. West of Gur’yev, the north coast of the Caspian Sea has a series of low spurs projecting southwest, separated by shallow bays. This irregular submerging coastline grades into the marshy islands and protrusions of the Volga delta, across the Russian border, which runs down one of the eastern distributary channels to the sea. In the northern Caspian region prevailing southeasterly winds produce strong wave action and occasional storms surges, which flood these low-lying areas.

References Ignatov YI, Kaplin PA, Lukyanova SA, Solovieva GD (1993) Evolution of the Caspian Sea coasts under conditions of sea level rise: model for coastal change under increasing “Greenhouse Effect”. J Coastal Res 9(1):50–57 Kaplin PA, Selivanov AO (1999) Sea level changes in Russia and coastal evolution: past, present and future (in Russian). GEOS, Moscow Rychagov GI (1997) Pleistocene history of the Caspian Sea (in Russian). Moscow University Press, Moscow Rychagov GI (2002) The Caspian Sea: past, present and future (in Russian). Zhyvopisnaya Rossia 3:39–42 Selivanov AO (1997) Coastal modifications of the Caspian Sea and other Central Asian lakes as natural models for coastal responses to the global sea level rise. Bollettino di Geofisica 37(1):114–129 Svitoch AA, Selivanov AO, Yanina TA (1998) Palaeogeographic Events in  the Ponto-Caspian and Mediterranean Basins. Moscow State University Svitovch AA, Yanina TA (1997) Quaternary sediments of the Caspian Sea (in Russian). Lomonosov Moscow State University Publication, Moscow

11.4 Russian Caspian Coast

Svetlana  Lukyanova

1.  Introduction In the Caspian Sea region, Russia possesses the northwest coast between the Volga and Samur deltas (>Fig. 11.4.1), a coast­­line of about 1,500 km. In terms of morphostructure, the northern part of this territory extends from the Scythian-Turan platform across the wide Terek-Caspian marginal trough, to the Great Caucasus megan­ticlinorium (>Fig. 11.4.2). These morphostructural features form the low-lying and flat morphology of the northern half of the  Russian coast, and mountains close to the sea form the southern half. This region has a dry, temperate continental climate. There are about 270 sunny days a year, and precipitation is usually not above 340–430 mm/year. Average annual air temperatures are about 12.5°C, and long-term summer temperatures range between 21 and 25°C. Steppe and forest-steppe vegetation predominates on this coast. The coast borders the shallowest northern and middle parts of the Caspian Sea, which determines specific features of the hydrological climate. Tide ranges are small, not more than several centimetres, but the coast is exposed to the prevailing strong southeasterly winds, which can provoke storm surges of up to 2–3 m, especially in the northern part of the coast. The southeast winds produce steep waves with a recurrence of 30% and heights of 2–3 m (a maximum height of 4.5 m having been recorded near Makhachkala). The stormiest sea is along the coast between Makhachkala and Derbent (with the number of stormy days numbering up to 89 in a year), where the influence of northeastern and eastern swells, with recurrence of 21–24%, is superimposed on the effect of storm waves. The predominance of longshore sediment migrations to the north on most parts of this coast is related to the wave climate; however, south of the Volga delta, the river discharge and alongshore wind-generated surge currents produce sediment transport to the south. The main sources of sediment in the coastal zone are the rivers, especially the Volga River, but also the Samur, Sulak, Terek, Shura-Ozen, Ulluchaij, and a few other rivers. On­­shore  movement of biogenic (shelly) sediment from the sea ­bottom is also very significant in some places. After ­implementation of hydro-engineering projects on most of the large rivers in

⊡⊡ Fig. 11.4.1 The Russian coast of the Caspian Sea. (Courtesy O.K. Leontiev.)

the ­second half of the twentieth century, and following active withdrawal of the river water for irrigation, the yield

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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11.4  

Russian Caspian Coast

⊡⊡ Fig. 11.4.2 Geological sketch map of the Caspian Sea coast: 1 – Quaternary deposits, 2 – Tertiary sedimentary rocks in folded zones, 3 – Tertiary rocks of platform zones, 4 – Mesozoic deposits, 5 – granites and g ­ ranodiorites, 6 – mud volcanoes, 7 – boundaries of structural zones. Structures zones: I-Russian ancient platform. II-Scythian-Turan young platform. III-TerekCaspian marginal trough. IV-Great Caucasus ­meganticlinorium. V-Kura-Araks depression. VI-West Turkmen marginal trough. VII-Talish ­anticlinorium. VIII-South Caspian basin. From Leontiev 1985.

of fluvial sediment to the coast has decreased considerably. Consequently, in some places there was erosion of sedimentary coasts even when sea level was falling. Long-term fluctuations in the sea level have played an important role in the development of the Caspian Sea coasts in general. These fluctuations have occurred repeatedly through the Pleistocene. Most investigators relate these sea

level oscillations to climatic factors. The Caspian Sea is a closed water system, sensitive to changes of water budget, with variations in fluvial discharge, precipitation, and evaporation substantially determined by climatic conditions in the drainage area and over the sea. The importance of climatic factors is confirmed by a correlation between the variations in the annual volume of Volga River discharge, which forms 70–80% of the total river discharge into the sea, and changes in the Caspian Sea level in the twentieth century. Early in the twentieth century, the Caspian sea level had only minor oscillations (an average long-term position of −26 m). A sharp fall in the sea level (by 1.8 m) occurred between 1929 and 1940. A number of subsequent smaller falls resulted in the emergence of extensive areas of the sea floor, in the appearance of inactive cliffs, and in the formation of two low and flat marine terraces around the perimetre of the Caspian Sea. The sea level declined to its lowest position (−29.02 m) in 1977, but since 1978 began to rise again, and by 1995 reached −26.6 m, a rise of almost 2.5 m. This has caused a general retreat of the coastline, flooding and waterlogging of agricultural fields, intensification of storm surge effects, and revival of wave erosion on previously abandoned cliffs. The cities of Lagan, Sulak, Makhachkala, Caspiysk, Derbent, and a number of smaller settlements have suffered from this sea level rise, but the major geomorphologic consequence is the wide development of lagoons and barriers on the flat young marine terraces. The low gradients (about 0.005 m) of the nearshore land that came under the sea influence resulted in formation of the narrow (about 200–300 m) and extensive (several kilometres in length) depositional barriers, backed by shallow lagoons. A slight lowering of sea level (by 0.54 m) since 1996 may be an episode that briefly interrupts an ongoing marine transgression, but it could mark the beginning of a new regression phase in the natural history of the Caspian Sea. Now (for a period of almost 8 years) the sea level stands at about −27.0 ± 0.2 m. This is following by lagoon draining and a slight widening of the young barriers on their seaward side.

2.  Russian Caspian Coastline Geomorphologically the Caspian coast of Russia is subdivided into four large sectors: the Volga delta, which is almost wholly included in the structure of the Astra­­khan area; the low-lying coast of Kalmykia; the low-lying ­Terek-Sulak plain of Northern Daghestan; and the narrow coastal plain with a staircase of marine terraces

Russian Caspian Coast

­ ordering Middle and Southern Daghestan. These areas b differ not only in the type of coasts, but also in their response to the modern sea level rise. The changes seen on the Caspian coasts during the sea level rise can be taken as a natural model for forecasting coastal changes under the acceleration of the global sea level rise that is  expected in the ­twenty-first century (Ignatov et al. 1993). The Volga River delta has a coastline of about 660 km (with the inclusion of islands). It is a compound (multilobe) delta on the northern section of the Russian coast of the Caspian Sea. Its surface lies between −22 m and −27 m, representing a low-lying Holocene and modern delta plain, with residual hills of protruding Upper Pleistocene ridge-and-hollow topography of the so-called Baer’s mound complex. The smaller units of the delta plain (channels, islands, oxbow lakes, natural levees, and coastal forms) formed in the course of the Holocene evolution of this delta, and make a uniform landscape. Baer’s mounds are Late Quaternary features widespread in the lowland bordering the northern Caspian region, particularly in the area of the Volga River delta. Morphologically, they are subdued parallel ridges, 5–22 m high, 0.5–10 km long, and 50–600 m wide, with sparse vegetation cover (Zenkovich and Popov 1980). Intervening swales are inundated by major floods of the Volga River. Baer’s mounds have a base of chocolate clays, and include layered loams with ­freshwater and brackish fossils and a fine sand cover. The origin of these mounds was first discussed by Karl Baer (1792–1876). For a long time they were regarded as ­aeolian features, but now the accepted view is that they formed in the sea (Svitoch and Klyuvitkina 2006) or on the wide lagoon floor and were modified by subaerial processes (such as wind and river action) when the Caspian Sea level was low in the Late Pleistocene. Others have suggested that they are of deltaic origin and similar to cheniers (Selivanov 1997). The transition from the delta plain to the gently sloping sea floor is gradual, through a wide transition zone of sandy mudflats. These mudflats are covered by hydrophytic vegetation, and exposed to seasonal flooding by river water and occasional flooding by seawater during storm surges, which reach a height of more than 2 m here. This transition zone is known as the kultuk zone because it is dominated by sediment deposition at the edge of the delta and in kultuks (bays between river mouths). There are low-lying sandy islands, large and small, as well as peninsulas, marshes, and banks dissected by the branching river channels and covered by dense reed vegetation. The marsh islands formed during the sea regression reached 60–80 sq. km. The modern kultuk zone is up

 11.4

to 40 km wide, narrowing from east to west, and is very dynamic, with a complicated interaction of fluvial (river channel processes, water, and sediment river discharge), marine (winds, waves, and currents) and biogenic (vegetation) factors. The Astrakhan State Nature Reserve (founded in 1919) is situated in this unique coastal zone. It has an area of 72,500 ha and includes both land and shallow sea areas that are a favourable habitat for various plants and birds. The various interacting processes cause constant changes along the delta coastline, with fluvial deposition processes having prevailed in the twentieth century. The river sediment discharge led to progradation of the outer part of the delta, especially during sea level lowering between 1929 and 1977. The delta front advanced by up to 20 km during the period between 1929 and 1950. The modern sea level rise, which began in 1978, reduced the rate of delta growth. In 2001, when the sea level was −27.20 m, the very wide shallow nearshore area (width 45–50 km, slopes less than 0.0002 m, depths slowly increasing from 0.5 to 2–3 m) weakened wave action and inhibited marine erosion on the submerging delta margin. Fluvial deposition continued, prograding sectors of the delta coast, especially in the semi-enclosed bays (kultuks) where most of the Volga sediment yield accumulates. In the early 1980s the average annual advance of the western part of the delta was 46.7 m, while that of the eastern part was 55.4 m (Rusakov 1989). Almost twenty years of sea level rise resulted only in the deepening of offshore water and partial modification of the system of nearshore sedimentary islands and some erosion of these. If the sea level reaches the threshold of −26.5 m, the erosion of the delta coastline will increase. The coast of Kalmykia extends for ­about 120 km from the Volga delta to Kizlyar Bay. It consists of extensive Pleistocene and Holocene marine plains, the surface of which is strongly modified by aeolian processes and covered with steppe vegetation. The younger (Holocene) plain, which is near the Volga delta, has parallel ridges (Baer’s mounds), (>Fig. 11.4.3) separated by wetland swales that form long, narrow inlets during major flooding of the Volga River or storm surges from the Caspian Sea. Low-lying marine terraces, which were formed during the sharp lowering of the sea level during the period 1929–1940, extend along the coast. Strong southeasterly winds generate powerful surges rising up to 1.5–2 m (at most, 3 m) above calm sea level – phenomena which are the main shore-forming factors here. As a result, the coast is fringed by a 6–8 km wide sand and mudflat that is subject to storm surge flooding and overgrown with dense reed vegetation. Gently sloping marshy plains extend

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⊡⊡ Fig. 11.4.3 A cliff cut into one of the Baer’s mounds on the northern Kalmykian coast. (Courtesy E. Badyukova.)

along the entire Kalmykia coast. It is difficult to identify the coastline precisely, because of the dense reed vegetation, the very low transverse gradients of the coastal plain and nearshore slope, and the variability of the sea level due to wind conditions. During the phase of falling sea level in the first half and especially in the middle of the twentieth century, this low-lying coast prograded rapidly; reed vegetation trapping surge-driven sediment promoted the deposition to some degree. The coastline advanced by at least 9 km (e.g. in the north of Kizlyar Bay) between 1948 and 1977. Analyses of maps and space imagery taken at various times have shown that the influence of the modern sea level rise along the Kalmykian coast increases from north to south as the transverse coastal zone gradient steepens (Kravtsova and Lukyanova 2000). This influence is minimal on the northern half of the coast adjoining the vast shoal of the Volga prodelta. The sediment deposition which was typical for the preceding period of sea level lowering still continues here (especially in small bays). In the southern section of the Kalmykian coast, south of Lagan, where the transverse slopes are a little steeper (but still only about 0.0005 m), sea level rise has affected a wide (up to 10 km) coastal fringe. The sand and mudflat zone has shifted landward, and its seaward margin has receded by up to 200 m/year as the result of submergence by the sea. There is also some wave modification of the upper part of the offshore zone due to the small increase of ­nearshore depths. As a result, a series of young beach ridges has formed along the water edge.

The Terek-Sulak plain extends south from Kizlyar Bay, and consists of a complex of Pleistocene and Holocene deltas built by the Terek River and the Sulak River. There are delta plains at different levels, numerous river arms frequently changing their channels, a series of natural levees, and also sets of natural and artificial reservoirs. The low, gently sloping northern part of the modern Terek delta plain extends to the southern coast of Kizlyar Bay. This bay is permanently filled up by thin deposits coming in by way of wind-induced surges from the shallow offshore zone, with river discharge, and also by an outlying branch of the Volga inflow. Deposition has been promoted by dense reed vegetation along the bay shores with a wide mudflat. The latter is composed of dark, silty sands with an abundance of vegetative remnants. In places, this mudflat is backed by a small (up to 0.5 m) cliff, cut by the most frequent storm surges. During the last thirty years of the Caspian Sea regression (1948–1978), the southern coast of Kizlyar Bay accreted up to 12 km as the result of passive emergence of the bay bottom due to sea level lowering and alluvial-marine deposition (>Fig. 11.4.4). To the south of Kizlyar Bay, the Terek-Sulak delta plain is bordered by several Holocene (Neo-Caspian) marine terraces. The present outlines of the coast in this area are determined by several large depositional features. From the north to the south, these include the Bryanskaya, Suyutkina, and Agrakhanskaya spits, and, further south, the ancient beak-like Sulak delta. The Bryanskaya and Suyutkina spits developed sequentially in periods of Holocene sea level rise and intensive

Russian Caspian Coast

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⊡⊡ Fig. 11.4.4 Emergence of Kizlyar Bay in the nineteenth and twentieth centuries: 1 – bay coastlines, 2 – Bryanskaya Spit, a Holocene feature. (Courtesy Lomonosov Moscow State University.)

wave erosion of the adjacent Terek delta. Eroded material drifted northward by the prevailing southeasterly waves nourished the growing spits. Evolution of these spits also continued in the mid-twentieth century, even when the sea level was falling, and was generally related to erosion of separate coastal sites under conditions of increasing sediment deficiency in the coastal zone. This deficit was a result of a gradual southward shift of the main Terek River outflow into Agrakhan Bay, and later (in 1973) by its ­artificial diversion through the Agrakhan Spit to the sea. Wave erosion of the proximal sites of the Bryanskaya and Suyutkina spits was already evident in the 1950s, and storm surges promoted this erosion to no small degree. Active cliffs, 1.5–3 m high, have receded here by ­150–200 m over a quarter of a century. The Agrakhan Peninsula is the largest depositional feature on the west coast of the Caspian Sea. The spit is extended in a submeridional direction to more than 50 km and has a width of 4–10 km. It originated at the beginning of the Holocene, on a nearshore shoal off the Terek delta, as a sandy barrier island backed by a shallow  lagoon  – the present Agrakhan Bay. Subsequently, under the influence of the intensive sediment discharge from the Sulak River, the barrier has joined the mainland and has formed a typical spit, with a series of diverging beach ridges and several islands of shelly sand in its distal part. The spit is capped by active dunes.

After the artificial relocation of the lower Sulak River segment in 1957, which diverted the main outflow to the southern part of its Holocene delta, sediment supply to the outer marine side of the Agrakhan Spit was drastically reduced, causing intensive erosion of its proximal part, even under conditions of sea level lowering. Over 25 years, a cliff 4–5 m high was cut into the sandy deposits of the barrier spit. The rate of coastal erosion was as much as 10–12 m/year. The modern sea level rise has accelerated this process. Diversion of the main Sulak River flow to the south has also had an effect on the evolution of the Holocene delta, which had previously actively accreted, growing to the north under the influence of strong southeast waves. During the previous century, the delta prograded by 9 km and underwent complicated evolution. After 1957 a young delta began to form rapidly at the new river mouth, which carried almost the whole volume of fluvial sediment discharge, and, by 1963, it extended up to 300–400 m into the sea. The old delta now receives very little fluvial sediment, and its outer edge has been cut back by up to 3 m/year by wave erosion. The Caspian Sea level rise that began in 1978 affected various types of coastlines differently. Low-lying marshes, especially within Kizlyar Bay, have been extensively submerged by the sea. Wave erosion has accelerated on the old cliffs along the southern parts of the Bryanskaya, Suyutkina, and Agrakhan spits. A narrow, shelly sand

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­ arrier with a similarly narrow elongated lagoon has been b formed along almost the whole eastern coast of the Agrakhan Spit (>Fig. 11.4.5), while the distal end of the spit has been submerged and waterlogged, and, in some places, has changed its outline. The formation of this extended coastal barrier resulted in significant realignment of the eastern coast of the Agra­ khan Spit. Its smooth outline is broken only by the small prominence of the “pioneer” Terek delta, which began to form here after the main river outflow was diverted to this site in 1973. However, with the rise in the sea level, this delta, as well as the similar young Sulak River delta, was eroded and partially flooded, and wide lagoons appeared on its surface. The old beak-like Sulak delta has undergone great modification as the result of the sea level rise (>Fig. 11.4.6). Space imagery shows that during the period 1978–1995 almost the entire part of this old delta was covered by sea water, and a narrow beach ridge was formed on its flooded surface. The rather narrow (1–2 –15–20 km) coastal plain of Piedmont (Middle and Southern) Daghestan has welldefined Pleistocene and Holocene marine terraces, which typically form clear steps and are very well expressed in relief. A series of Holocene barriers extend along ancient lagoons such as Big and Small Turaly and Adjy. As a whole, the present coastal zone of this area generally has more steep transverse slopes (about 0.005–0.01 m) and receives

stronger wave action, which dominates shore processes here. During the mid- twentieth century lowering of the Caspian sea level, low-lying emerged marine terraces (1929–1940), fringed by wide (up to 150 m) sandy beaches, were formed along practically the entire coast of the Piedmont Daghestan. There were active dunes behind the beach, especially to the south of Izberbash Town. Processes of wave erosion weakened appreciably. However, in the later years of the regression (1960–1970s), coastal erosion became more active because of sediment deficiency in the coastal zone due to the reduction of river discharge (caused by hydrotechnical engineering in river valleys, water withdrawal for irrigation, rectification of river channels) and wave erosion downdrift from seaport structures. At present, coastal erosion has accelerated as the result of the modern sea level rise. The previously formed low terraces and beaches have been submerged, and the sea has advanced towards the older bluffs. The outer edge of the Neo-Caspian (Holocene) marine terrace is now intensively eroded north of Makhachkala (>Fig. 11.4.7), where a strip of land about 150 m wide has been lost, and the present cliff is up to 4 m high. Coastal erosion is most obvious at and near the large cities of this coast, such as Makhachkala, Caspiysk, and Derbent, where artificial structures have resulted in acceleration of coastal bluff erosion. Thus, the erosion of

⊡⊡ Fig. 11.4.5 Barrier-lagoon complex formed during the modern sea level rise along the eastern coast of the Agrakhan Spit. (Courtesy L. Nikiforov.)

Russian Caspian Coast

⊡⊡ Fig. 11.4.6 Recent changes on the Sulak delta. 1 – coastline in 1978, 2 – coastline in 1991, 3 – submerged parts of the delta, 4 – lagoons, 5 – barriers formed on the submerged delta surface, 6 – land, 7 – Caspian Sea. (Courtesy V.I. Kravtsova.)

a coastal bluff 4–5 m high, cut in the Holocene marine terrace on which Caspiysk City is situated, resulted from the trapping of northward alongshore sediment trans­ port by port breakwaters. The lower erosion of the lee­ ward coast is apparent here, at a distance of about 3 km. Quays, ladders, and artificial embankments erected earlier in the city have changed natural slopes and also promoted

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activation of coastal erosion at a modern rate of 6–8 m/ year, threatening urban structures. A similar situation has developed near Makhachkala, where the rate of coastal retreat has reached 10–12 m/year. Coastal protection works have not always been effective. The pile-protected structures built at Caspiysk City have been almost completely destroyed by the sea. An attempt to protect a sanatorium by stone riprap and reinforced concrete blocks represented a temporary measure against the concentration of wave energy on the coastal prominence formed here. This building was 150 m from the sea at the moment of its construction, but is now under threat of wave destruction. Bedrock outcrops occur on the narrow coastal plain fronting the Large Caucasus mountains and on the underwater coastal slope, especially along the anticlinal tectonic structures near the sea. The resistant bedrock outcrops produce benches of various types, stony capes, and ridges, which are the usual units of the underwater relief of this region, such as near Izberbash Town. When the sea level was falling, the rocky ridges impeded wave erosion. During the sea level rise, water depths above the ridges increased, but they still play a part in resisting erosion (>Fig. 11.4.8). Artificial structures play the same role in some places. On low-lying sedimentary parts of the Piedmont Daghestan coast, the falling sea level led to the formation of wide sandy or shelly beaches backed by active dunes, but a barrier beach up to 2 m high has developed during the modern sea level rise. This feature encloses a shallow lagoon (with a depth of less than 1.5 m), which is filled by storm wave overwash across the ridge and the upward movement of groundwater table accompanying the sea level rise. This type of coast is seen, for example, to the south of Caspiysk, between the Satun and Buynak capes and along the Izberbash-Derbent segment (>Fig. 11.4.9). The newly-formed barrier is a single sedimentary feature extending up to several kilometres, simple in outline, and typically composed of coarse, shelly sand. The primary indicator of transgressive coastal development is the gradual landward movement of the barrier towards the adjacent lagoon. Repeated profiling of the barrier has shown that the rate of its retreat is about 30 m/year (Ignatov et al. 1993), but a slight fall in the sea level over the last few years has briefly slowed this process. The modern marine transgression has indirectly led to weakening of the aeolian processes. Upward movement of the groundwater table accompanying the sea level rise has increased wetness in dune areas, which have been overgrown with vegetation, while the narrow (10–15 m)

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⊡⊡ Fig. 11.4.7 Erosion of the Holocene marine terrace north of Makhachkala. (Courtesy L. Nikiforov.)

⊡⊡ Fig. 11.4.8 Outcrops of densely lithified Middle Pleistocene c­ onglomerate forming Satun Cape to the south of Caspiysk City. (Courtesy L. Nikiforov.)

transgressive beaches have no time to dry to a degree sufficient to permit any appreciable sand movement by the wind. As a result, active dunes have practically disappeared on this coast. Overall, processes of the coastline retreat have prevailed on the Russian coast during the 20 years of the

modern Caspian sea level rise. Various mechanisms may have been responsible for this retreat: the wave erosion of  sea cliffs, the shifting of the storm surge mudflats, or the  landward movement of the lagoon-barrier systems. Coastal accretion persists only in the far north, especially along the shores of the Volga delta.

Russian Caspian Coast

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⊡⊡ Fig. 11.4.9 The modern barrier to the south of Isberbash Town. (Courtesy L. Nikiforov.)

References Ignatov YI, Kaplin PA, Lukyanova SA, Solovieva GD (1993) Evolution of the Caspian Sea coasts under conditions of sea-level rise: model for coastal change under increasing “Greenhouse Effect”. J Coastal Res 9:104–111 Kravtsova VI, Lukyanova SA (2000) Studies of recent changes in the Caspian coastal zone of Russia based on aerial and space imagery. J Coastal Res 16:196–206 Leontiev OK (1985) Caspian USSR. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 481–486

Rusakov GV (1989) Main delta-forming factors and formation of the marine edge relief of the Volga delta and its prodelta under conditions of the Caspian Sea level changes and economic action. Moscow State University Selivanov AO (1997) Coastal dunes in Russia under the rising sea level: present and future management aspects. In: Ovesen CH (ed) Coastal dunes - management, protection and research. Skagen, Denmark, pp 96–104 Svitoch AA, Klyuvitina TS (2006) The Baer’s mounds of Lower Volga. Press MGU (in Russian), Moscow Zenkovich VP, Popov BA (eds) (1980) Marine geomorphology: Coastal Zone. Moscow

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11.5 Azerbaijan

Pavel Kaplin · Andrei Selivanov · Svetlana Lukyanova

1.  Introduction The Azerbaijan coastline (>Fig.  11.5.1) is about 800 km long, at the southeastern end of the Caucasus Ranges (Leontiev 1985). It has a dry subtropical climate. The Caspian Sea has fluctuations of up to 3 m level, because of wind action and barometric pressure. Tide ranges are small, only a few centimetres. Prevailing easterly winds produce strong wave action and occasional storm surges. Beach sediment has been supplied partly by rivers, and shelly material and oolites have been swept in from the sea floor, with ⊡⊡ Fig. 11.5.1 Structure of Eastern Azerbaijan: 1 – Peri-Caspian trough and coastal lowland; 2 – Great Caucasus; 3 – Durbar; 4 – Apsheron and Gobustan Plain; 5 – Kura Lowland; 6 – Talysh Ridge. (Courtesy K.A. Veliev.)

local accessions from eroding cliffs. Longshore drifting is northward to the north of the Apsheron Peninsula, but south of this peninsula longshore drifting becomes southward as NE winds and waves become more effective than waves generated by SE winds over a diminishing fetch. The morphology of the Caspian Sea is related to geology, with the Derbent depression (780 m deep) and the South Caspian depression (1,022 m deep) corresponding with structural troughs. Separated from the oceans in Lower Pliocene times, the Caspian Sea has shown fluctuations related to climatic changes within the bordering   drainage basins. During the Quaternary there were repeated glaciations of the Russian Plain and the Caucasus Mountains, and intervening deglaciations. The level of the Caspian Sea consequently rose and fell in the Pleistocene, attaining a level 50 m higher than at present in the Middle Pleistocene and falling 50 m below present level in the Late Pleistocene. Marine terraces correspond with stillstands during these fluctuations. In Azerbaijan tectonic movements associated with the uplift of the Caucasus Mountains have raised the oldest coastlines 200–300 m above present level. During Holocene times a sea level rise, the NeoCaspian marine transgression, culminated in the eighteenth century, when it attained 6 m above its present level. Since then there has been a substantial lowering, with a sharp fall between 1930 and 1976 (>Fig.  11.5.2), which reduced the area of the Caspian Sea by more than 40,000 sq km. During this sea level fall there was emergence of land margins and shoals became islands; there was shoreward drifting of shelly calcareous sand from the sea floor on to beaches, spits and barriers (Leontiev et al. 1977). The sea declined to its lowest level (−29.02 m) in 1977, but then began to rise again, and by 1995 reached −26.6 m, a rise of almost 2.5 m. This has caused a general retreat of the coastline, flooding and waterlogging of agricultural land, intensification of storm surge effects and revival of wave erosion on previously abandoned cliffs. Since 1978, the proportion of eroded coastline has increased from 20 to 55%, and sea floor sediment, primarily shelly debris, are increasingly dominating beaches. Beach erosion is also partly due to the reduction of fluvial sediment supply ­following the damming of rivers and diversions into

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 11.5.2 Changes in the level of the Caspian Sea from 1930, as indicated on the Baku tide gauge. Sea level fell intermittently until 1977, then rose fairly rapidly.

i­ rrigation canals. Coastal towns have suffered from the sea level rise, but the main geomorphologic consequence is the wide development of transgressive lagoons and barriers. The low gradients of the coast resulted in marine submergence generating narrow (about 200–300 m) and extensive (several kilometres in length) depositional barriers fronting shallow lagoons, all transgressive as sea level rise proceeded. The implications of these changes for other coasts in the expected global sea level rise have been much discussed (Ignatov et al. 1993; Selivanov 1997; Kaplin and Selivanov 1995; Rychagov 2002).

2.  Coastline of Azerbaijan The Samur River on the border between Azerbaijan and Russia was the primary source of sediment to the northern coast of Azerbaijan, but diversion of river flow to irrigation channels halved this sediment supply and led to erosion of the delta. Sediment is still drifting south, derived from the eroding delta as well as the diminished fluvial supply, but alongshore drifting diminishes southward, as does the extent and thickness of Holocene depositional formations. South from the Russian border the main (Great) Caucasian Mountain Range approaches the coastline from the northwest. Numerous sub-parallel rivers flow down the northeastern slopes and across the coastal plain to the sea. The main geological structure in the north is SamurDivichi depression, which is a foredeep filled with Upper

Pliocene and Quaternary sediment more than 500 m thick. They are mostly deltaic-alluvium from the Samur and other numerous small rivers. Alluvial-deltaic gravels outcrop in a cliff near the Nabran village and to the north of it. The Samur-Divichi plain gradually declines eastward to the sea. The coastal plain narrows southward from the SamurDivichi plain to no more than 10–12 km in the Kilyasi district and to only a few hundred metres near Sumgait City. It is backed by a series of uplifted Neogene and Quaternary terraces rising to 320 m. The coast has sandy beaches, backed by extensive dunes. South from the Samur delta the coastline is smooth, a confluent delta plain fringed by beaches and dunes, and the Divichi flows into the Akzybir Lagoon, which lies landward of a sandy barrier. This lagoon became a marshland when the level of the Caspian Sea fell (Leontiev et al. 1977), but it is reviving as a result of the modern transgression and at present is a reed swamp. The Kilyazi protrusion is an anticlinal ridge of Jurassic rocks fringed by beach ridges. It grew by up to 40 m/year during the sea level fall (1930–1976), but growth slackened and ceased during the ensuing sea level rise. The outer part of the Kilyazi Cape is composed of bedrocks (flysch rocks of the anticline core) and has typical wavecut bench with numerous small ridges formed by selected erosion of the flysch. In the bay to the south the Sumgait River flows from the declining southeastern end of the Caucasus Mountains down into a bay beside the coastal town of Sumgait.



The Apsheron Peninsula is a hilly protrusion 60 km eastward, with alternating headlands of Neogene rocks and broad open bays fringed by beaches now much reduced due to submergence and erosion. It culminates in a narrow spur with northward growing spits on its eastern shore. Artem and Zhiloy are outlying islands. Erosion and submergence have resulted in the building of extensive protective structures, including sea walls and breakwaters. Baku is a major city and port on the southern side of the Apsheron Peninsula, a projection of the Caucasus Ranges into the Caspian Sea with bays on synclines bet­ ween rocky promontories on anticlines in Upper Cre­ taceous formations. It has a stylish promenade looking out to the bay. The port has many breakwaters. Buzovnaya beach is up to 350 m wide and has been maintained even during the modern phase of sea level rise by artificial nourishment. On the other hand the Shivova spit has been much reduced in recent years. South of Baku the narrow Gobustan coastal plain extends southward with capes of Late Pliocene limestone alternating with wide bays containing broad beaches. Beyond Primorsy and a small delta at Duvannyy it begins to widen to the Muganskaya Plain, which is drained by the Kura River and its large tributary, the Araks. There are networks of irrigation canals. The Kura has built a protruding delta, and longshore drifting is now southward, indicated by the direction of growth of the Kura spit in front of Kirov Bay (Zaliv Kirova). This spit was initiated as a ­barrier beach formed from sediment brought down the Kura

⊡⊡ Fig. 11.5.3 Erosion on the Caspian coast south of Lenkoran, showing broken coastal structures now up to 30 m offshore. (Courtesy E. Ignatov.)

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11.5

River. But about 200 years ago the river breached it and began building its modern delta projecting seaward. In recent years parts of the delta coastline have been cut back by erosion. Southward drifting of fluvial sediment continued to nourish the barrier beach, prolonging it as a spit, but after the construction of a dam on the Kura River the sediment supply diminished, and the seaward shore of the Kura spit began to erode while accretion continued at its southern end. It still separates Kirov Bay from the Caspian Sea during low sea level periods, but becomes a chain of shelly sand islands when sea level is high. During the 1929–1976 lowering of sea level, the Kirov Bay became marshy in the lee of the spit, but it is now again largely submerged, with a second, inner spit extending south to Narimanabad. South of the Kura lowland the Talysh Mountains (Little Caucasus) come in from the northwest, with rivers draining their northeastern slopes feeding sandy sediment to beaches. These also receive shelly calcareous sand from the sea floor, but erosion has become prevalent with the sea level rise during the past two decades. Beach erosion is severe near Lenkoran, and the coastline is retreating by up to 25 m/year. During a major storm in October 1990 the coastline retreated by up to 6 m (UNEP 1997). The  foothills of the Talysh Plain are fringed by a coastal plain, which narrows from 10 km at Lenkoran southward to less than 2 km on the Iranian border at Astara. It is undergoing erosion (>Fig. 11.5.3) and sea walls have been built in several places (>Fig. 11.5.4).

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⊡⊡ Fig. 11.5.4 Beach erosion in the northern part of Astara City threatens the railway station. (Courtesy E. Ignatov.)

References Ignatov YI, Kaplin PA, Lukyanova SA, Solovieva GD (1993) Evolution of the Caspian Sea coasts under conditions of sea level rise: model for coastal change under increasing “Greenhouse Effect”. J Coastal Res 9:50–57 Kaplin PA, Selivanov AO (1995) Recent coastal evolution of the Caspian Sea as a natural model for coastal responses to the possible accelerated global sea level rise. Mar Geol 124:161–175 Leontiev OK (1985) Caspian USSR. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 481–486

Leontiev OK, Maiev EG, Rychagov GI (1977) Geomorphology of coasts and floor of the Caspian Sea (in Russian). Moscow University Press, Moscow, p 209 Rychagov GI (2002) The Caspian Sea: past, present and future (in Russian). Zhyvopisnaya Rossia 3:39–42 Selivanov AO (1997) Coastal modifications of the Caspian Sea and other Central Asian lakes as natural models for coastal responses to the global sea level rise. Bollettino di Geofisica 37:114–129 UNEP (United Nations Environmental Programme) (1997) Consequences of climate changes in the Caspian Sea in the Caspian Sea Region. UNEP, Geneva

11.6 Aral Sea

Eric Bird

Introduction The Aral Sea (>Fig. 11.6.1) lies 500 km east of the Caspian. Its northern half is in Kazakhstan, its southern half in Uzbekistan. Most atlas maps show its 1960 outline, when it was the world’s fourth largest inland water area and stood 53 m above world sea level. It occupied a tectonic depression that formed early in Pleistocene times, bordered by a high cliff in Cretaceous sandstones (probably a fault scarp) on the eastern side of the Vost Chink Ustyurta Range, which rises 250 m above sea level. In 1960, the Aral Sea had a volume of about 1,100 cu km and an area of about 68,000 sq km. It extended about 400 km from north to south and 250 km east to west, and had an average depth of 21 m rising to a maximum of 68 m. ⊡⊡ Fig. 11.6.1 The Aral Sea, showing the 1960 coastline and the extent of the North and South Aral Seas in 2006. Compiled from various sources. (Courtesy Geostudies.)

In the NW, the coast was indented between hilly promontories of Cretaceous rock, but the rest was lowlying, a coastal plain of sandy Quaternary sediment, with the large silty deltas of the Syr-Darya in the NE and the Amu-Darya on the south coast. There were numerous small islands. The climate is arid, with cold winters and hot summers with a little rain. Prevailing winds are NE in winter and W to SW in summer. 90% of the fresh water flowing into the Aral Sea in 1960 came from the two rivers, Amu-Darya and Syr-Darya, and inflow was balanced by evaporation. Salinity of much of the Aral Sea was about 10 ppt (parts per thousand), rising to over 100 ppt in shallow bays, where evaporite deposits (mainly sodium sulfate and sodium chloride) formed and were harvested. The Amu-Darya (classical Oxus) rises in the Pamir Mountains and is fed by glaciers. It flows NW to the south coast of the Aral Sea, where by 1960 it had built a vast delta with many abandoned distributary channels. Its flow was double that of the Syr-Darya, which rises in the Kirghiz Mountains and flows through the Fergana Basin and out NW to the NE coast of the Aral Sea, where it had built another large delta below Kazalinsk. Since 1960, the Aral Sea has been greatly reduced. By 2006, its level had fallen 26 m and it had been reduced to less than half its 1960 volume and area and split in two (North and South Aral Seas). 38,000 sq km of the former Aral Sea floor had emerged as bordering land, and the coastline had advanced by up to 155 km to border the shrunken seas. Numerous sub-parallel beach ridges (contraction ridges) mark stages in the lowering of sea level, probably separated as the result of occasional storms or minor oscillations of sea level. Salinity has increased and the sea water is polluted with pesticides and fertilizers. The 1960 harbours have been abandoned, and steamships lie stranded on the emerged sea floor. The fishery has declined, and biodiversity has been reduced. It was expected that the South Aral Sea would have disappeared by 2020 (Jones 2003). Reduction of the Aral Sea followed the large-scale diversion of water from the two rivers into reservoirs and canals for the irrigation of cotton and rice, which began in the 1930s. The Amu-Darya and Syr-Darya deltas have

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_11.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Aral Sea

emerged, and the diminished outflow runs down into the lowered Aral Seas. Most of the delta distributary channels are dry. A small outflow continues on the eastern side of the Amu-Darya delta and there is intermittent discharge from the Syr-Darya. Outflow channels are slightly incised across the emerged sea floor. In 1987, an isthmus emerged to separate a smaller North Aral Sea from a larger South Aral Sea. A dam (Dike Kokaral) was completed in 2005 to keep them separate, and divert the flow from the SyrDarya into the North Aral Sea. Since then, the North Aral Sea has risen sharply while the South Aral Sea continues to dry out. The present coastline of the South Aral Sea borders a sandy and salty plain, subject to dust storms, but the rise in the North Aral Sea is a transgression similar to that seen in the Caspian Sea.

There is evidence of earlier fluctuations in the level of the Aral Sea, related to climatic changes rather than human impacts. The post-1960 emergence has revealed archaeological evidence of fourteenth century burial sites, indicating that the water level was lower then. There is also evidence of former tree growth in areas that were submerged in 1960. It may well be that the post-1960 lowering was not entirely due to the diversion of river water into irrigation areas.

Reference Jones N (2003) South Aral gone in 15 years. New Scientist, 2404 (19 July 2003): see also New Scientist (8 April 2006)

12.0 North Africa – Editorial Introduction

Africa stands on a tectonic plate that is advancing north towards Europe. North Africa has a number of structural units that influence the outlines of the coast, generally with an E–W trend (Orme 2005). The Nile delta dominates the Egyptian coast (> Egypt (Mediterranean)), where beaches and dunes are of quartzose sand supplied from river distributaries, but these pass westward into beaches and beach and dune ridges of calcareous sand, including Pleistocene dune calcarenites. Beach rock is commonly exposed. In western Egypt and >Libya, the coast is low-lying with cliffs cut in Tertiary and Quaternary formations, including dune calcarenite. The Jebel el Akbar range in Cyrenaica (E. Libya) forms steep rocky coasts with terraces at several levels, then the low-lying coast between Benghazi and Cape Misratah, intersecting the Sirte Basin (filled with Tertiary sediment) is fringed by dune-capped barriers, lagoons, marshes and supratidal flats, notably Sabkha az Tayurglia. The Jabal Tarhuna ridge comes to the coast at Al Khums, but west of Tripoli similar features fringe the Gefara Plain. Mean spring tide range is generally less than 0.5 m, increasing to 1.8 m in the Gulf of Gabes. The prevailing NW winds generate swell and waves from that direction, producing an eastward longshore drift. In > Tunisia, the coast turns northward in the Gulf of  Gabes, and becomes hilly and indented as ridges of Tertiary rock protrude. These are the eastern end of ranges produced by Alpine folding, one of which runs out to Cape Bon. The Medjerda River has built a delta into the Gulf of Tunis. At Bizerte, the coast turns westward to run

along the strike of the folded Tertiary formations, and at Cape Blanc there are cliffs cut in Eocene limestones capped by Pleistocene dune calcarenite. The Algerian coast (> Algeria) follows the trend of the  Tell, where Alpine folding produced a succession of  ranges, with many sectors that are steep and cliffed. The outcrops include Pre-Cambrian crystalline basement, Mesozoic limestones and sandstones, a variety of Tertiary sedimentary and volcanic formations and Quaternary deposits, notably dune calcarenites. Beaches of sand and shingle are fed by rivers such as the gravelly Moulouya. There are sectors of swampy lowland, and basins of Qua­ ternary sediment. Oran has high limestone cliffs, and to the west the coast follows the northern flanks of the Rif Mountains into > Morocco, curving out to Ceuta and the Strait of Gibraltar. Neotectonic movements have caused uplift and depression of coastal sectors, and earthquakes indicate continuing activity near the margin of the African plate. The harbour of ancient Carthage, near Tunis, is now beneath the sea, and stairways of marine terraces extending along uplifted sectors of the Algerian coast. Mean spring tide ranges are small (Fig. 12.1.1). There are four lagoons on the northern section of the 250-km-long delta shoreline, increasing in area from Port Said to Alexandria: Lake Mariut, Lake Idku, Lake Burullus, and Lake Manzala. All four lakes are shallow and brackish, fed both by freshwater canals flowing from the main two tributaries of the Nile and by the sea. Lake Manzala, through which the canal passes, has many small sandy and marshy islands, and is separated from the sea by a narrow, curved sandy barrier. Southeast of Damietta there are many arcuate, deserted, narrow barriers, representing former coastlines. These run out to a cuspate point at the mouth of the Damietta branch of the Nile. Beyond is a broad bay, backed by deltaic plains through which the Bar Shibin distributary formerly flowed. A lobate point, Cape Burullus, is followed to the west by Lake Burullus, another marshy lagoon separated from the sea by a narrow, sandy barrier, with a gap at El Burg near the eastern end. Since the building of the Aswan High Dam in the early 1960s, this barrier has been migrating landward. Another cuspate point marks the outlet of the Rosetta branch of the River Nile, and to the west is rounded Abu Qir Bay, backed by Lake Idku behind another narrow barrier, in the form of spits on either side of an entrance from the sea. Abu-Qir Bay is the site where Napoleon’s fleet was defeated by Nelson in 1898. At the western end of

Abu-Qir Bay, a small promontory runs out to a pier, and to the southwest is the seafront of Alexandria. The present delta is the product of deposition that began in the Upper Miocene, some 10 million years ago. Nile sediment was generally deposited in shallow water, in an area where the Earth’s crust has subsided nearly 3 km, partly because of the weight of the accumulating sediment load (deltaic cone) and partly because of tectonic movements associated with the creation of the present Mediterranean. Fine-grained Nile sediment originate from the alumina and iron-rich soil of Ethiopia, and have particular depositional features distinct from those of other soils of the Mediterranean region (Manohar 1981). The main clay minerals are montmorillonite (smektite), illite, and kaolinite, which form most of the fluvial load and are spread on the outskirts of the Nile cone and in the south-eastern corner of the Levantine Basin. The coarse fraction, which is mostly composed of quartz, also originates in Ethiopia and in Sudan, mainly from eroded Nubian sandstones. These sands used to arrive mainly during the June to September Nile floods. Additional sand originates in the Eastern and Western Desert of Sudan and the numerous Egypt wadis that flow very irregularly. These sediment averaged about 80–160 million tonnes per year before the completion of the Aswan High Dam in 1964, but were then reduced to insignificant quantities. The reduction in sediment supply began with the building of the Delta Barrages, completed in 1881, and the Aswan Dam completed in 1902, but major changes followed closure of the High Dam in 1964. These include severe erosion along the delta coastline and the penetration of saline water from the sea into river mouths, lagoons, and groundwater. The coastline of the Nile Delta has undergone severe changes during the past century (Stanley 1990; Stanley and Warne 1998). Erosion has occurred at several locations along the coast (>Fig. 12.1.2), notably at El Gamil between Port Said and Ras el Barr, on either side of the Damietta outlet at Ras el Barr, at El Burg, on either side of the Rosetta outlet, at Idku, and at Alexandria (El-Ashry 1979). El Burg, near the Lake Burullus outlet, is one of the most exposed parts of the delta. The coastline is straight for a long stretch along the barrier enclosing Lake Burullus, and is oriented obliquely to the prevailing northwesterly wave direction, so that sand drifts alongshore to the eastnortheast. The slope of the nearshore is also relatively steep, allowing the waves to reach the shore without being completely adjusted or dissipating their energy. As a result, beach erosion has occurred along this section for a long time, and serious property damage has been incurred.

Egypt (Mediterranean)

12.1

895

⊡⊡ Fig. 12.1.1 The Nile Delta, showing some of the former distributaries. (Courtesy Geostudies.)

⊡⊡ Fig. 12.1.2 The coastline of the Nile Delta, indicating sectors of rapid erosion. (Courtesy Geostudies.) El Burg

Cape Burullus N

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12.1

Egypt (Mediterranean)

Beach erosion has been rapid on the promontories at the mouths of the Rosetta and Damietta branches of the Nile. On the western side of the Rosetta mouth, the promontory has lost about 1,650 m of its length in 65 years, averaging about 29 m/year. According to Wassing (1964), the coastline retreated at 15 m/year during 1898–1926, 81 m/year during 1918–1926, 21 m/year during 1926–1944, and 35 m/year during 1944–1954. In comparison, the foreland at Ras el Barr has retreated by about 1,800 m between 1902 and 1960, with an average of 31 m per year. Beach erosion has been accompanied by the concentration of heavy minerals in shore sands (Frihy 2003). Whereas the ongoing coastal erosion to the west is due to wave and longshore impact, erosion to the eastern part of the delta is also due to subsidence (Stanley 1990). Between the eroded sectors, there has been deposition of sand supplied by the two remaining distributaries of the Nile and sand eroded from the delta coastline. The depositional features include beaches, spits, forelands, offshore bars, and barriers enclosing lagoons (El-Ashry 1979). There are up to three offshore bars, an inner break-point bar, which forms at a depth of about 2 m and migrates shoreward during the summer, when another bar forms at 2 m depth (Manohar 1981). Seaward, intermediate and lower bars form at depths of 3.0–3.5 m and 4.5–5.5 m during storm episodes. West of Alexandria the beach becomes a barrier enclosing Lake Maryut, and the coast curves round to the breakwater at Dukhaylah before trending west-southwest into the Persian Gulf (Warne and Stanley 1993). Here,  the coast is backed by parallel ridges of dune calcarenite separating sebkhas. The main direction of longshore drift can be seen in the asymmetrical bays and the erosion on the eastern sides of promontories. North of El Alamein the continental shelf becomes narrower, and the low-lying coast gives place to low dune calcarenite and limestone cliffs on headlands, as at Ras-el-Daba and Ras el Kenaysis, a cuspate cape beyond Kenaysis Bay. A ­pediment ­descending to the coast is incised by many wadis, and the Mersa Matruh lagoon is one of the small bays, bordered by

enclosing spits. The rising ground behind the coast includes a series of Pleistocene ridges representing ancient coastlines. Butzer (1960) identified and mapped eight such ridges, corresponding to at least eight marine transgressions above present mean sea level during the Pleistocene. The lower and younger seaward ridges have scrubby vegetation cover, which becomes sparser on the older and higher ridges in a drier setting landward. Between the ridges are swales, formerly elongated lagoons and now evaporite corridors (El-Asmar and Wood 2000). A long lobate coastline extends past Sidi Barrani into the Bay of Salum, where an elongated sebkha backs the coastal dune ridge. The small port of Salum shelters in a bay, with Beacon Point to the north, as the coastline swings northward across the Libyan border.

References Butzer KW (1960) On the Pleistocene shorelines of Arabs’ Gulf, Egypt. J Geol 68:626–637 El-Ashry MT (1979) Use of Apollo-Soyuz photographs in coastal studies. In: El-Baz F and Warner DM (eds) Apollo-Soyuz test project: summary science report, vol 11, Earth observations and photography, Washington, DC, pp 531–543 El-Asmar HM, Wood P (2000) Quaternary shoreline development: the northwestern coast of Egypt. Quat Sci Rev 19:1137–1149 Frihy E (2003) The Nile delta-Alexandria coast: vulnerability to sea-level rise, consequences and adaptation. Mitigation and Adaptation Strategies for Global Change 8:115–138 Frihy OE, Lotfy MF (1997) Shoreline changes and beach sand sorting along the northern Sinai coast of Egypt. Geo-Mar Lett 17:140–146 Manohar M (1981) Coastal processes at the Nile Delta coast. Shore Beach 49:8–15 Stanley DJ (1990) Recent subsidence and northeast tilting of the Nile delta, Egypt. Mar Geol 94:147–154 Stanley DJ, Warne AG (1998) Nile delta in its destruction phase. J Coastal Res 14:794–825 Warne AG, Stanley DJ (1993) Late Quaternary evolution of the northwest Nile Delta and adjacent coast in the Alexandria Region, Egypt. J Coastal Res 9:F 26–64 Wassing (1964) Coastal Engineering Problems in the Delta Region of United Arab Republic. Egyptian Department of Ports and Light­ houses, Alexandria

12.2  Libya

Maurice Schwartz

1.  Introduction The coast of Libya (>Fig.  12.2.1 A,B), along the northern border of the Sahara desert, extends for 1,850 km, from Egypt in the east to Tunisia in the west. A coastal highway runs between Benghazi and Tripoli, the two major ports of the country. Oil transshipment ports are located along the coast in proximity to interior oil fields. There is a fishing industry based on sardines, tuna, and sponges, and agriculture is centred in a narrow (up to 10 km wide) coastal plain in the west, where olive, almond, and citrus trees are grown. Temperatures in the coastal region range from 12°C to 16°C in January to 24°C to 28°C in July. Rainfall is sparse and variable, with warm tropical continental winds blowing offshore in January and onshore winds in July. A Mediterranean climate prevails in the east and west, with a desert in the central part of the coast. Much of the coast is bordered by Mediterranean steppe vegetation, and the soil is the reddish-brown to the grey and greyishbrown of a subdesert zone, and is often saline. The wave environment along the Libyan coast is that of a protected sea, with less than a 10% frequency of waves of 1.6 m height or more during at least two quarters of the year. The dominant swell is from the northwest (Orme 2005). Tides are semidiurnal in the western half and mixed in the east, and spring tide range is less than 2 m. Average annual rainfall in the coastal zone is less than 50 mm. The geology of the coastal region (Burollet et al. 1971) is briefly as follows: (a) the Pre-cambrian African shield is composed of gneiss, schist, quartzite, and granite; (b) Palaeozoic deposits are sedimentary, and have undergone tectonic vertical movements and the formation of basins; (c) block faulting and flexures form the present structures (e.g., the Gulf of Sirte), where Cenozoic sediment have accumulated. Tectonically, the region is associated with the Ionian Plate in a northwesterly direction of plate motion (Orme 2005). The present-day surficial geology consists of Neogene continental formations in the east and undifferentiated Quaternary deposits in the west. The Barce plateau in the eastern coastal region of Cyrenacia has a maximum elevation of 900 m, reaches 48 km in width, and is 240 km long along the coast. The major sea cliffs in Libya are developed

where the plateau ends in steep slopes along the seaward margin. Beaches are mainly sandy, composed partly of terrigenous sand eroded from cliffs and carried down wadis during occasional phases of runoff and partly of calcareous sediment from sea floor organisms and the Pleistocene aeolianites, which fringe parts of the coast and extend beneath the sea.

2.  The Libyan Coast The eastern end of the coast, near the Libya-Egypt border north of Umm Sa’ad, is sandy on the western side of the Gulf of Sollum. Near Sawani al Mallahah the coastline swings westward, and has cliffs on the coastal slopes, with white dunes in a low-lying sector. Several wadis descend from a plateau hinterland to the coast. Tobruk is a port in the lee of a headland, and to the west there are more white dunes behind sandy beaches in the Gulf of Bomba. Wadis descend from the Al Jebal al Akhdar limestone mountain range, where there are rocky cliffs and narrow emerged terraces. The coastline again swings west at the Ra’s at Tun cape, with cliffs cut into the lower slopes of the mountain range, extending past Derna to Ras Hilal and almost to Tukrah. The archaeological site of Apollonia is at Marsa Susah, 100 km west of Derna. From the Tukrah region to Benghazi the coast is lowlying and sandy. Benghazi is a major city and oil-exporting port. The coast swings southward alongside the Gulf of Sirte, and the beaches become barriers backed by lagoons and sebkhas (dry areas, except during exceptionally high tides or occasional rainy periods), continuing round to Al-Uqaylah. From here to Sidra there are dunes behind the beach, and a sandy desert hinterland; this is one of the most arid parts of the Mediterranean Sea coast. To the west there are some sectors of low cliff alternating with dunes behind sandy beaches. Wadi Thamit ends in a usually dry saline inlet at the coast. Beach ridges and swales along the 100 km coast that curves northward from Al Bu’aryat are backed by sebkhas, including the wide Sabkhat Tawurgha. At Cape Misratah the coastline swings westward, past the towns of Misratah and

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_12.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Libya

⊡⊡Fig. 12.2.1 The coast of Libya: A. geography; B. geomorphology. (Courtesy Geostudies.)

Zlitan, to Al Khums (Horns), with the archaeological site of Lebdah (Leptis Magna). Inland, The Jabal Nafusah limestone range runs inland behind the Gefara Plain to the west. The coast has intermittent sandy sectors between steep hilly stretches along to the port of Tripoli (Tarabulus), built on a coastal salient, with a harbour protected by large breakwaters. There are sandy beaches on either side of the harbour, and Wadi al Majinin reaches the sea on the western side. Pleistocene dune calcarenite borders the coast in the broad bay that extends past Az Zawiyah, with sectors of low cliff and some shore platforms exposed at low tide.

From Az Zawiyah to the Tunisian border the coast is sandy, backed in places by sebkhas. Westward drifting has built a spit sheltering a sebkha at Abu Kammash.

References Burollet PF, Magnier P, Manderscheid G (1971) La Libye. Tectonique de L’Afrique, UNESCO, Paris, pp 409–416 Orme AR (2005) Africa, coastal morphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 9–21

12.3 Tunisia

1.  Introduction The coastline of Tunisia is about 1,300 km long, including the shores of coastal lagoons and islands. The region has a dry Mediterranean climate, generally low to moderate wave energy, and microtidal conditions. There are marked contrasts between the eastern and northern coasts on either side of Cape Bon.

2.  The Tunisian Coastline The eastern coast of Tunisia extends from the Libyan border to the Cap Bon peninsula and borders the eastern Mediterranean. Semiarid conditions dominate in the Gulf of Gates, where the mean annual rainfall is less than 200 mm, but rainfall increases northward. The continental shelf is broad, more than 200 km wide at the latitude of the Kerkennah Islands. Only local winds, blowing from between the east and north in spring and summer and from the west in autumn and winter, affect the coastline, which is characterised by low energy. Waves mainly come from the southeast in the Gulf of Gates, and from the northeast north of the Kerkennah Islands. They generally die out before reaching the shore because of the shallowness of the wide continental shelf. Submarine meadows of phanerogams (Cymodocea nodosa and Posidonia oceanica), which extend offshore from depths of 1–30 m, also help diminish wave energy, and also promote sand deposition. On Jerba Island accretion that is taking place at the tip of the Ras Rmel Spit is the counterpart of severe erosion of the long beach fringing the northeastern coast of the island (Miossec and Paskoff 1979). Erosion probably started after the sixteenth century, as the recurved spit of Ras Rmel did not then exist, according to a map dating back to 1560. Recently, coastline recession has been dramatically increased by unwise human intervention related to the tourist boom that has affected the island during the last two decades. There has been sand mining on the beach

for building purposes, and in many places the foredune has been completely destroyed to build hotels very near to the shore. The removal of dead leaves of Posidonia, which accumulate on the shore and inhibit sunbathing, increases the effects of wave action, which is also enhanced by the degradation of the infralittoral meadows damaged by sewage. The Kerkennah Islands protect the segment of coast between Gates and Sfax, but during exceptional winter storms wave erosion may occur. Generally, there is weak longshore drifting southward. Tides are important in the Gulf of Gates, with a maximum range reaching 1.8 m, and tidal currents of up to two knots are especially noticeable in the narrows surrounding Jerba and Kerkennah Islands, where they account for the occurrence of the so-called submarine wadis, which are meandering channels excavated in the shallow continental shelf. The geomorphic features of the eastern coast of Tunisia are explained by the subdued topography of the hinterland together with the low-energy marine environment. Cliffs are small and infrequent. Sandy beaches predominate, particularly around the Gulf of Hammamet, in the vicinity of Mahdia and Gabes, in eastern Jerba, and near Zarzis. In the last two areas, beach rock is occasionally found. It appears that the sand of the beaches fringing the Gulf of Gabes came from the sea floor during the Versilian transgression; at present there is no significant source of sand supply from wadis or eroding cliffs on this coast. Near Gabes, beaches are highly polluted by wastes from industrial plants erected during the last decade. Muddy tidal flats extend south of Sfax. Elsewhere, low rocky platforms cut in Tyrrhenian marine or aeolian sandstones (Paskoff and Sanlaville 1979) are found south of Zarzis, in eastern Jerba, and on the Kerkennah Islands. The Monastir area (>Fig.  12.3.1) on the east coast of Tunisia is well known for its Tyrrhenian deposits rich in Strombus bubonius. The term Monastirian, although no longer in general use, was for some time widely employed to designate a marine Quaternary stage in the Medi­ terranean. The name was first applied to differentiate a

Edited version of a chapter by R. Paskoff in The World’s Coastline (1985: 523–530). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_12.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

12. 3

Tunisia

⊡⊡ Fig. 12.3.1 Geomorphic sketch of the Monastir area. 1. marine cliff cut in Neogene rocks; 2. low rocky coast; 3. sandy shoreline; 4. erosional scarp of tectonic origin; 5. normal faults; 6. strike-slip movement; 7. coastal barrier of Eutyrrhenian age; 8. wave-cut platform veneered by Eutyrrhenian deposits. (Courtesy Geostudies.)

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lower 15 m marine level, extending south to Monastir and separated by a cliff running along the wadi Tefla, from a higher 30 m marine level attributed to the Tyrrhenian. Both levels include Strombus in their surficial deposits. It  is now established that there is only one wide wavecut  platform at Monastir. This platform, veneered by Eutyrrhenian deposits (Rejiche Formation), has been faulted and tilted. A hinge fault, corresponding to the wadi Tefla thalweg and facing southeastward, was erroneously interpreted as a former sea cliff; it separates two compartments that have been unequally uplifted. Another major post-Tyrrhenian fault is the SkanesKhniss fault, which runs near the foot of the scarp limiting the platform of Monastir westward and crosses the former coastal barrier of Khniss, which runs near the foot of the scarp limiting the platform of Monastir westward and indicates the position of the coastline at the maximum of

the Eutyrrhenian transgression. Field evidence shows that this fault has both a horizontal and a vertical component. Important post-Tyrrhenian tectonic events in the surroundings of Monastir are well documented. These events controlled the present coastal morphology of the area, characterised eastward by a low rocky shore developed in Eutyrrhenian sandstones and northward by retreating cliffs cut in soft Neogene rocks (Paskoff and Sanlavllle 1983). The Gulf of Tunis marks a transition between the eastern and northern coasts of Tunisia. It is relatively protected from the northwesterly swell by its northeastward orientation. Large embayments are controlled by a more open continental topography. Long sandy beaches correspond to deltaic plains built since the end of the Versilian transgression by the only two perennial streams of Tunisia, the Miliane River and the Medjerda River, both prone to heavy floods, especially the latter. The delta of the Medjerda River (mean annual ­discharge: 30 m3/s) is a good example of a delta in a Med­ iterranean environment (>Fig. 12.3.2). Its development is fairly well known from geological and geomorphological studies, corroborated by archaeological data. The delta has formed during the last 5,000–6,000 years by the filling, from north to south, of a former bay drowned by the Versilian marine transgression. Human activity played an important role in its evolution. Deforestation in ancient times considerably increased the solid load carried by the river, and public works (artificial levees, man-made cutoffs, drainage and diversion canals) completed during the present century have interfered with natural processes. For instance, during a heavy flood in March 1973, with the water flow reaching 3,500 m3/s (a figure higher than the estimated maximum 100-year rate of discharge), the Medjerda River shifted its lowermost course. The natural channel was abandoned and the entire flow now passes through an artificial canal, originally designed to evacuate the excess of water during floods and discharge it directly to the sea (Paskoff 1978). As a result, the recurved, spit which has developed at the former mouth since the end of  the last century, is presently suffering erosion at its root,  extending its tip and migrating westward. Trans­ formations of the coast on both sides of the new mouth, located about 10 km south of the former one, are also ­considerable. A cuspate foreland shows rapid extension despite sediment trapping behind dams constructed on the Medjerda River and some of its tributaries. Between 1977 and 1980 delta progradation added about 35 ha of new land to the coast. West of the Medjerda delta the building, in 1975, of the outer harbour of Ghar el Melh, on the barrier that isolates

12. 3

Tunisia

⊡⊡ Fig. 12.3.2 The Medjerda Delta. (Courtesy Geostudies.) N

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a lagoon from the sea, has resulted in accretion on the updrift (western) side of the jetties protecting the entrance of the dock and erosion on the downdrift sector, which has been depleted of drifting beach material to such a degree that the barrier has been cut off. The northern coast from Bizerta to the Algerian border relates to the western Mediterranean. The continental shelf is narrow, less than 10 km wide. The coast is rainy (annual average rainfall: 600 mm) and windy. The prevailing winds (in both strength and frequency) blow from the  northwest; consequently, the coastline is exposed to northwesterly swell. Winter storm waves are generally several metres high. Longshore drifting is eastward. The tide is semidiurnal, and its average range is small, about 0.2 m. The coastline truncates the highlands of northern

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Tunisia (Mogod and Kroumiri mountains) and, because of high wave energy, rugged cliffs and bluffs, cut principally in Oligocene sandstones, represent the main geomorphic features. Only short ephemeral streams reach the sea, and there are few beaches west of Tabarka. However, an abundant sand supply derived from eroding cliffs and the regional climatic conditions (strong winds, marked summer dryness) generates extensive mobile dunes (about 300 sq. km) of quartzose sand, showing evidence of southeastward movement and transgressing over the hinterland topography. Almost the entire Tunisian coastline has been eroded, and numerous valuable coastal sites of archaeological ­interest have been destroyed by wave action. A slight sea level rise during the historical period partly accounts for

901

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12. 3

Tunisia

this ­situation. Beach erosion is a serious problem in Tunisia. The tourist industry, mainly based on sea and sun, has been a major priority of the authorities since 1960 and has become the second largest source of foreign currency for the country. In fact, with the exception of a few restricted sectors of progradation, notably near the mouth of the Medjerda River, the sandy coasts are threatened by erosion. Rapidly retreating beaches are found in and around Tunis, particularly on Jerba Island, as indicated previously.

References Miossec JM, Paskoff R (1979) Evolution des plages et aménagements ­touristiques à Jerba (Tunisie): le cas du littoral nord-est de l’île. Meditérranée 1–2:99–106 Paskoff R (1978) Evolution de l’embouchure de la Medjerda (Tunisie). Photo-Interprétation 5:1–23 Paskoff R, Sanlaville P (1983) Les côtes de la Tunisie, variations du niveau  marin depuis le Tyrrhénien. Collection Maison de l’Orient Méditerranéen, 14 Lyon, 192pp

12.4 Algeria

M. Larid

1.  Introduction The Algerian coastline is about 1,100 km long,  from 2° 10' W to 8°30' E in longitude and from 35° to 37° N in latitude (>Fig. 12.4.1). The coastline is backed by chains of mountains with slopes plunging abruptly into deep

water and intervening sandy low coasts, which ­represent about 20% of the coastline (Larid 2001). The climate of the Algerian coast is Mediterranean, the western zone being drier than the eastern. Rain-bearing winds from the northwest influence the distri­bution of vegetation, soil erosion, and the flow of water and sediment down wadis.

⊡⊡ Fig. 12.4.1 Geomorphology of the Algerian coastline. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_12.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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12.4

Algeria

Onshore winds come mainly from the northwest round to the northeast. These produce waves that arrive mainly from the north-northwest in winter, averaging 2–3  m in height, and reaching 5–6 m during storms; in summer the waves come from the north-northeast. The north-northeasterly swells (20–40°) prevailing in summer are less strong, averaging 0.5–1.0 m in height and rarely reaching 3 m. Tide ranges along the Algerian coast do not exceed 0.5 m, and tidal currents are weak. The cool Atlantic current, less saline than the Mediterranean Sea, comes through the Strait of Gibraltar and along the Algerian coast, and produces eddies in the larger bays (Gulf d’Arzew and Gulf de Bougie). It may sift sea floor sediment, but is of little importance in coastal geomorphology. The main features of the Algerian coast are related to the crystalline massifs and to the Mesozoic calcareous or schistose massifs (Bourcart and Glangeaud 1954). There are headlands where abrupt cliffs, 100 m in high, descend steeply beneath the sea. Between these are bays and gulfs cut into younger, horizontal, or gently dipping rock formations, with low coasts, sometimes rocky and indented (cut in Pleistocene dune calcarenite), or sandy beaches with Holocene dunes. Cliffs are cut in various rock formations. Some are actively receding, as in the Pleistocene dune calcarenite, which forms rugged cliffs with solution weathering (lapies) fronted by shore platforms; others are plunging cliffs, steep rocky slopes that have been little modified by marine erosion at present sea level. Some abandoned cliffs stand back from the coast. There are also low, rocky coasts, and shore platforms with algal terraces. Rivers deliver sand, silt, and clay to the coast and the continental shelf. Winter rainfall can be heavy, causing erosion in the river catchments, where the slopes are abrupt and the degraded vegetation cover insufficient to prevent rapid runoff. The sediment yield can be considerable. The average annual sediment load delivered to the sea from the wadis of Algeria has been calculated at 40–60 million tonnes. In recent decades there has been a reduction in rainfall and consequently in runoff and sediment yield. Beaches of sand and gravel are usually found at or near the mouths of the wadis, and are mainly composed of terrigenous sediment brought down by wadi streams to the coast and distributed by wave action along the adjacent shores. Other beaches have been derived from the erosion of cliffs and rocky shores, particularly in Pleistocene dune calcarenite, and some sandy sediment (including shelly material) has been washed in from the sea floor and deposited on beaches. Spits are found, particularly at the

mouths of wadis, where the stream outflow has been too weak to prevent sand and gravel drifting alongshore to form spits that deflect stream mouths. These spits are variable in configuration because of the interaction of stream outflow and longshore drifting. Some spits become temporary barriers that seal off wadi mouths as lagoons (guelta in Arabic) or swamps until the next stream flood breaches them. In addition to spits there are tombolos, where sand deposition has attached islands to the mainland, as at Sidi Fredj (west of Algiers, where longshore drift converges from opposite directions) and Cap Falcon near Oran. They are not currently very active, and were probably built during the early Holocene. Small tombolos have also formed in the lee of nearshore breakwaters, as on beaches at Moretti and Sidi Frèdj. Many beaches are eroded during winter storms. At the mouth of Isser Wadi (Kabylia) the beach is cut back by up to 40 m in winter, and waves reach the Pleistocene dunes. During summer the beaches are restored, with deposition of large quantities of fine sand. Many (about 70–85%) of the beaches have been cut back by erosion in recent decades, partly because of stormier conditions in coastal waters, partly because of the quarrying of beach sand for building and road-making, and partly because runoff and sediment supply from streams has been reduced by drier conditions and barrage construction. The beach east of Bedjaia retreated about 300 m between 1980 and 1998. Coastal dunes are mainly found near large wadi mouths where the relief is low and the sea floor shelves gently, as at the Mazafran, Isser, Seybous, and El Kebir wadis. Some dunes are derived from Pleistocene dune calcarenite formations eroded by marine action, the sandy products of which are blown by the wind. There are actively-forming dunes at Isser and Jijel, and dunes being eroded by the sea at Annaba. The discharge of silt-laden water from certain wadis has permitted the formation of salt marshes and mudflats near their mouths, particularly where they are sheltered by sand spits, as in the La Macta swamps near Oran and at the mouths of the Seybouse and El Kebir wadis in the east of the country.

2.  The Algerian Coastline West of the Tunisian border the coast crosses a zone of strongly folded Tertiary rocks, where mountains with a north-south geological strike reach the sea. Beyond Cape Rosa is an embayment backed by the Fetzara Plain, where

Algeria

Quaternary sedimentary formations occupy the El Kebir and Sebouse wadis. The coast is low-lying, and fringed with Pleistocene dune calcarenite and Holocene dunes. Beyond Annaba there are steep coastal slopes with basal cliffs cut in the metamorphic rocks of Mount Edough, and similar profiles are cut in Oligocene sandstone from the Seraidi beach to Chetaibi. From Chetaibi to the Cape de Fer the coast is steep and cliffy, cut in volcanic rocks. Next comes the big depression of Lake Fetzara, separating the coastal massif of the Edough from the inland Mounts of Guelma. The coast is low and swampy, with dunes, between Cape de Fer and the Djebel Filfilah. West of Djebel Filfilah (near Skikda) the crystalline massif of the Collo-Kabylia forms a high and abrupt façade, rising to 950 m a few kilometres from the shore. To  the west the Neogenic basin of Jijel opens to a relatively low coast, with abandoned cliffs cut in Tertiary and Quaternary deposits. The volcanic outcrop of Cavallo and the limestone range of Babors form a cliffy coast, called the Corniche Kabyle. From the Bay of Bejaia to Dellys, the Great Kabylia is bordered to the north by Tertiary and Mesozoic sandstones and schists in mountains that border the sea in the Bay of Bejaia. This bay corresponds to the tectonic valley of the Wadi Soummam. Between Dellys and Cape Matifou there are large accumulations of sand, especially between Cape Tamentefoust and Cape Djinet, and a narrower sandy beach fronts abandoned cliffs between Cape Bengut and Cape Djinet. The Bay of Algiers lies between Cape Matifou (a metamorphic rock outcrop) and the Bouzareah massif. This low-lying coast is occupied by the suburbs of Algiers, which spread from the Bay of Zemmouri in the east to the Bay of Bou Ismail in the west. Between the metamorphic ­massif (crystalline schists) of Bouzareah and the Jurassic

12.4

uplands of Chenoua the coast is rocky, cut into Holoce­ne Recent (Villafranchian) formations, and slightly elevated in the Bay of Bou Ismail. From Tipaza (west of Algiers) to Mostaganem, the Dahra coast ranges (Tertiary sandstone) and the schistose massif (flysch) of Beni Menacer descend to the sea from a hundred metres. Between the sandstone plateau of Mosta­ ganem and the Djebel Orous the Gulf of Arzew is backed by the extensive low and swampy plain of La Macta, where alluvium has been deposited by the Wadis Mekerra and El Hammam. In the Gulf of Arzew the coast is cliffy, cut in Quaternary formations, with the Carbon Cape and Cape de l’Aiguille cut into the calcareous Djebel Orous. The calcareous massif of the Oran Region rises more than 400 m above the sea, except around the tombolo of Cape Falcon and the nearby low, sandy coast. From Cape Sigale to Cape Figalo volcanic andesites form steep cliffs, while from Cape Figalo to Beni Saf the coast is cut into Villafranchian and Quaternary continental sediments. The Traras Mountains, consisting of schists, quartzites, and crystalline limestones, form a steep cliffy coast about 200 m high, plunging into water more than 50 m deep. The western part of the Algerian coast is low-lying, with sand and shingle beaches derived from sediment carried down by the Moulouya Wadi (from Morocco) that drifted eastward.

References Bourcart J, Glangeaud L (1954) Morpho-tectonique de la marge continentale nordafricaine. Bulletin de la Societé Géologique de la France 4:751–772 Larid M (2001) Littoral Algérien: état et perspectives. Etude réalisée pour le compte du ministère de l’aménagement du territoire et de l’environnement, ISMAL - Algiers

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12.5 Morocco

Anja Scheffers

1.  Introduction The Moroccan coast is about 2,500 km in length, extending more than 350 km from the Algerian border westward along the Mediterranean Sea to the Strait of  Gibraltar, then southwest along the Atlantic Ocean (>Fig.  12.5.1). The Mediterranean coasts border the Rif Mountains and the Tangiers Peninsula. Lengthened considerably (by another 1,400 km) since the annexation of the Western Sahara in 1976, the Atlantic coasts between the Rharb and the Anti-Atlas (which is, truly speaking, Atlantic Morocco) have been the subject of many studies. The coasts of Saharan Morocco are still little known. Morocco has a warm temperate maritime climate, becoming arid in the south, with northeast trade winds dominant in summer and northwest winds in winter. These produce seasonal variations in waves from these directions. Rabat has mean monthly temperatures of 12.9°C in January, 22.2°C in July, and an average annual rainfall of 564 mm. Mean spring tide ranges are 2.9 m at Casablanca and 1.9 m at Dakhla. While the Mediterranean coast has waves generated by local winds, the Atlantic coast also receives a northwesterly ocean swell.

2.  The Mediterranean Coast The Mediterranean coast of Morocco is mostly mountainous, with beaches of sand and pebbles along wide embayments for about 25% of its length. The Moulouya delta has been built just west of the Algerian border, and the coast to the west is bordered by sandy beaches (>Fig. 12.5.2). Near Nador a 25-km spit has grown westward to shelter a large lagoon (Sebkha Bou Areq) on the flanks of the narrow promontory that, 20 km north to the Cap des Trois Fourches, juts into the sea. The coast to the west has a succession of cliffy promontories between bays with sand and shingle beaches, the gravel prominent at the mouths of valleys such as Wadi Kerte. The bay of Al Hoceima is sandy between cliffed

headlands. The hinterland rises to the Rif Mountains, which trend roughly parallel to the coastline, and descend to cliffs that are up to 300 m high in the Bokkoya. Between Al Hoceima and Ceuta is a broad arcuate bay, the coast following the structural trend as it curves out to Cape Almina. The coastline then turns west along the southern side of the Strait of Gibraltar, past Ceuta to Cape Spartel, and cuts across the structural trend of the mountain ranges. In general, headlands with rocky cliffs correspond to the folds, and bays with beaches to the synclines. The Mediterranean coast of Morocco borders a geosynclinal alpine chain. Variations in the levels of emerged Pleistocene and Holocene terraces (>Fig. 12.5.3) indicate that there have been neotectonic movements: transverse warping and faulting (Brückner 1986). The mountain ranges of the Rif also influence the northern section of the Atlantic coast, from Cape Spartel to Larache. The coast, which is gently curved, cuts across strongly folded and complex thrust sheets of friable rocks, with Cretaceous sandstones and marly sands being predominant (Orme 2005). Beaches are extensive, shaped by ocean waves, between sectors of low cliff cut into the Cretaceous sandstones and marls, and also Pleistocene dune calcarenites (aeolian sandstones). The coastline has been smoothed by long-continued erosion and deposition, with adjustment to prevailing northwest swell waves. Much of the Atlantic coast of Morocco has low cliffs (often less than 10 m high) in the north, rising to more than 30 m towards the south. Beaches extend for more than 20 km, about 30% of the coastline (>Fig. 12.5.4). As the continental shelf is not very wide (Summerhayes et al. 1971), the coastline, which acquired its primary characteristics as early as the Moghrebian (Middle-Upper Pliocene) transgression, has varied little during the Quater­nary changes of sea level. The resulting general layout, when seen in profile, appears composed of three main elements (>Fig. 12.5.5). There is an upper platform, here, about 20 km wide and usually above 100 m high. This is known as the Moghrebian Rasa (the term rasa is used in

Edited version of a chapter by A. Weisrock in The World’s Coastline (1985: 537–544). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_12.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

12.5    

Morocco

⊡⊡ Fig. 12.5.1 Structure and predominant coastal landforms of Morocco. (Courtesy Geostudies.) 10° W

MEDITERRANEAN SEA Rif Rharb

modeiro

OC

EA N

1

Meseta

Hoouz

2

as High Atl N 30°

s Sou Conary Islands

Anti

as

- Atl

LA

NT

IC

3

AT

908

T.C. Reguibat Arch

0

400 km

200

MOROCCO: STRUCTURE AND PREDOMINANT COASTAL LANDFORMS Coastal basins : 1: Safi-2: Atlantic Atlas

4

3: Western Sahara- 4: Mauritanian

cliffs & bluffs

sandy shorelines

Morocco  

⊡⊡ Fig. 12.5.2 The Moroccan Mediterranean coast. (Courtesy Geostudies.)

0

Ceuta Cap spartel

   12.5

Tanger

50 km

Cap des Trois Fourches Nador

Al Hoceima B.Soid

Bokkoyo

te

Sebkha Bou-Areg uy a

O.

O

.M

ou

2452

r Ke

lo

Larache

NORTHERN MOROCCO - PREDOMINANT COASTAL LANDFORMS

⊡⊡ Fig. 12.5.3 Coastal terraces of Beni Said (after Barathon 1978). (Courtesy Geostudies.)

Main ranges of the Rif

bluffs

Volcano

beaches

cliffs

high cliffs

Beni Saïd shales SE

Marine deposits

NW 4

60 m

Dune sandstones 3

Slope debris Calcareous crust 2

40

20

1 0

Steplike marine terraces of the Beni Saïd (after Barathon 1978)

Morocco to describe a platform, probably of marine ­origin). It has been folded, and culminates in the Atlantic Atlas at a height of more than 600 m. There is also a lower platform, which formed in several stages during the Quaternary rasa, and is up to 2 km wide and close to the present coastline. The most typical aspect of this lower platform (rasa) is the bench (oulja), which takes the shape of a downfold parallel to the coastline, sometimes enclosing a lagoon. This bench can also be made up of several step-like marine terraces. Between the upper and lower ­terraces is an abandoned cliff, which marks the landward limit of Quaternary marine transgressions. The strong Atlantic swell has cut cliffs in dune calcarenite, but

the large boulders (some more than 40 tonnes) found many metres high on the seaward calcarenite ridge and up to 100 m inland, often finely imbricated, are believed to be the result of tsunamis, most probably from the 1755 ad Lisbon earthquake (Mhammdi et al. 2008). There are variations in these coastal features, depending on the height of the abandoned cliff and its position in respect to the present coastline. There are also extensive coastal dunes, derived from former marine sands, which have developed during long summer droughts, with a strong prevailing northeast trade wind driving them towards the Atlantic coast. Some of these dunes have been cemented by precipitated carbonates into dune ­calcar­enite,

909

12.5    

Morocco

MOROCCAN ATLANTIC COAST GEOLOGY N

Quaternary Moghrebian sandstones Tertiary and Mesozoic sediment of the Rif

RIF 1000

N

20

0

A E

Mesozoic sediments

O

C

Newer volcanics Hercynian Caledonian

⊡⊡ Fig. 12.5.4 Geology and landforms of the Moroccan Atlantic coast. (Courtesy Geostudies.)

8°

RHARB

34°

Rabat

Casablanca

Granite

IC

Sidi Moussa

L

A

N

T

Oualidia Cap Cantin

M

T

Jarf el Yhoudi

E

S

E

T

A TADLA

A

A

910

T

L

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PREDOMINANT Cap Hadid Essaouira

ESSAOUIRA BASIN

COASTAL LANDFORMS HAOUZ 0

T. Ouftas

ATLANTIC ATLAS

HIGH

Moghrebian dead cliff Quaternary

Cap Rhir

SO

30°

100 km

50

„

„

Low cliff

US

High cliff Moghrebian rasa Oulja

Ifni

T I N S A A L T A O.D

ra

Sandy shorelines Coastal dunes Merja

OW

and there is evidence of five successive phases of dune calcarenite generation. Around 30° N chevrons occur as parabolic sand deposits on top of cliffs where no beach sand is available. The coast also shows variations that depend on hinterland topography. Low coasts are found along the gulfs that

occupy the subsided plains of the Rharb and the Sous; high coasts, though fairly rare, are found where the Meseta, the Atlantic Atlas, and the Anti-Atlas form three projections to the west (>Fig. 12.5.6). The longest beaches primarily border the plains of the Rharb (110 km) and of the Sous (60 km), and secondly the

Morocco  

⊡⊡ Fig. 12.5.5 Some landform associations of the Moroccan Atlantic coast: Rharb, Meseta, and Atlantic Atlas. (Courtesy Geostudies.)

   12.5

RHARB: South of Larache

50

merja 6

126

MESETA: Cape Hadid ouljo

10

Maghrebian

359

ATLANTIC ATLAS: Cape Rhir

46 20 10

MOROCCAN ATLANTIC COAST: SOME ASPECTS

region of Essaouira (20 km). These beaches are supplied with sand carried down the rivers by winter floods. Onshore northwest winter winds supply sand from these beaches to some very large coastal dune complexes, up to 80 m high (as in the Rharb), backed by marshes (merja) towards the interior. Beneath the high Pliocene Moghrebian platform (Sahel), the Meseta Coast is a typical oulja-coast, where the Palaeozoic basement and its covering strata outcrop only rarely as small monoclinal capes (skhour). The

oulja is sometimes bordered towards the ocean by a narrow beach, as is often the case between Casablanca and Cape Cantin, or it can end in a low cliff cut into the dune ­calcarenites. Bioerosion has produced visors and low tide shore platforms, as well as rugged pitted cliffs, particularly south of Rabat. Behind the foredune the northern marshes give place to lagoons (Oualidia) or tidal marshes (Sidi Moussa). The high cliffs from Cape Cantin to Jorf el Yhoudi are an exception; in this area, the oulja has disappeared and the abandoned cliff has

911

912

12.5    

Morocco

⊡⊡ Fig. 12.5.6 Holocene seif dunes and mobile sand dunes at Cape Sim, near Essaouira. The natural plant cover (a forest of Juniperus phoenicea) has virtually disappeared, except on the crusted seif dunes. Fixation of the barchan dunes is taking place, but sand blown by the NNE trade winds continues to supply the beach. (Courtesy P. Oliva.)

come to the coast, eroded either at its base or across its full height. This erosion of the oulja, or of the lower platform, occurs along the Atlantic Atlas, where low cliffs have been cut in the Cretaceous basement, in the Pleistocene dune calcarenites, or in Quaternary alluvium. The abandoned cliff comes close to the coast at Ouftas and merges with the modern cliff on Cape Rhir. Emerged beaches have been preserved on some of the headlands, notably Cape Rhir, indicating Quaternary stages that can be compared with the sequences at Casablanca and Rabat. To the souththe Anti-Atlas coast presents two main aspects. Along the Sous the dunes are still predominant, and low cliffs have been cut into thick calcareous sandstones. In the Ifni district there is a coast of crystalline rock overlain by Cambrian formations. which outcrop in rugged cliffs. Small beaches occupy coves beneath wide, stepped marine terraces.

3.  Western Sahara Stretching for more than 1,400 km to the south, the coast of Western Sahara (formerly Saraoui, Spanish Sahara, or Rio de Oro) has a general outline that is made up of long embayments between minor capes (>Fig. 12.5.7), and the

­ orphology does not change much along all this stretch m of coast. There are cliffs cut into Tertiary and Quaternary sandstones and limestones. As the trade winds drive Atlantic water offshore and cause cold water upwelling, this is an arid coastal region, with no permanent rivers. Offshore are the arid Canary Islands. Coastal outlines are influenced by northwesterly ocean swell that generates longshore drifting southward (Guilcher 1954), and there is also sand movement from NNE–SSW, produced by the trade winds, with some sand blown into the sea (if not to the Canary Islands). Towards Cap Juby the coast faces northward, and the hinterland has arrays of large seif dunes parallel to the trade wind, extending for long ­distances (65 km from Megriou to the Seguia el Hamra). Winds were in a similar direction in Plio-Pleistocene times, since they formed the old dunes and spits that date from the Moghrebian, which are now consolidated dune sandstones, often cut into low cliffs. The prominent fixed Afrafir coastal sand ridge runs from south of Cap Juby, across the mouth of the Seguia el Hamra wadi, and then passes inland. To the south of the Sebkha Amtal, the landscape of small crusted dunes is cut again by the ocean in the Aguerguer, where there are many sebkhas in the small slacks between the dunes. These sebkhas are saline areas subject to recurrent marine flooding. Vila Cisneros is a narrow promontory that runs southward, sheltering a

Morocco  

   12.5

⊡⊡ Fig. 12.5.7 Geology and coastal landforms of Western Sahara. (Courtesy Geostudies.) 10°

15°

O.D

ra

fra

25

aA

0

fir

Cap Juby

Dra

Km SCALE 1 cm = 50 km

Seguia e

Homra

Cap Bojador

25°

Ri

od

eO

ro

Villa Cisneros

Cap Barbas

Cöte rocheuse

Sandy coast

Cöte sableuse

Seb

kh

a

I

l d ji

Rocky coast

Pre-Cambrian border Palaeozoic border

Cordon d’ Aguerguer

Mesozoic Tartiory

Cap Blanc Port Etienne

shallow bay, Rio de Oro. The coast is low and sandy from here to Cape Barbas, about 200 km south. The Pliocene Moghrebian terrace is rarely more than 40 m above sea level, and cliffs capped by fossiliferous Moghrebian deposits cut into it on Cape Boujdour and between Cape Barbas and Dakhla are no higher than this.

Limites du Précambrien Limites du Primaire Secondaire Tertiaire

Quaternary

Quaternaire

Fixed dune ridges

Cordons de dunes fixées

The lower platform is less continuous than in northern Morocco, but there are isolated remains of ouljian beaches found there, as are remains of the sebkhas, which are ­associated with the last Nouakchottian coastline, formed during a marine transgression about 5,500 years ago (Ortlieb 1975).

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Morocco

References Barathon JJ (1978) Some aspects of the recent geomorphological evolution of the northeast Rif. (in French) Etude de certains milieux du Maroc, C.N.R.S., Montpellier 4:1–23 Brückner H (1986) Stratigraphy, Evolution and Age of Quaternary Marine Terraces in Morocco and Spain. Zeitschrift für Geomorphologie, NF, Supplementband 62:83–101 Guilcher A (1954) Dynamics and morphology of sandy coasts in Atlantic Africa. (in French) Information Géographique 1:57–68

Mhammdi N, Medina F, Kelletat D (2008) Isolated boulders along the Rabat coast: a possible tsunami-related origin? Science of Tsunami Hazards 21:17–30 Orme AR (2005) Africa, Coastal Geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, pp 9–21 Ortlieb L (1975) Research on the Plio-Quaternary formations of the West Saharan coast. (in French) Travaux et Documents, O.R.S.TO.M., Paris Summerhayes CP, Nutter AH, Tooms JS (1971) Geological structure and development of the continental margin of northwest Africa. Marine Geology 11:1–25

13.0 West Africa – Editorial Introduction

The African Plate began to separate from the South American Plate in Jurassic times, and the intervening area has since become the Atlantic Ocean. The West African coast has three main divisions: from the Strait of Gibraltar south to Cape Verde in Senegal; from Cape Verde to the mouth of the Congo; and from there to the Cape of Good Hope (Orme 2005). It is covered in > Mauritania, > Senegal and Gambia, > Guinea Bissau, > Republic of Guinea, > Sierra Leone, > Liberia, > Ivory Coast, > Ghana, > Togo and Benin, > Nigeria, > Cameroon and Equatorial Guinea, > Gabon, Congo, Cabinda and Zaïre, > Angola and > Namibia. The Atlantic coast of Morocco runs across the axes of  the Atlas Mountains, which consist of several parallel ranges produced by Alpine folding and trending northeast to southwest. There are narrow coastal plains between ridges that protrude in headlands. Rivers draining the mountains supply sediment to the coast, particularly ­ during the rainy winter season. The mean spring tide range diminishes south from Casablanca (2.9 m) to Dakar (1.3 m) and the coast is exposed to northwesterly swell, which results in southward longshore drifting. To the south, in > Mauritania, the coast is low and cliffed, and Saharan dunes are driven towards it by northeast trade winds. The desert extends to the coast and includes Late Tertiary sandstones, dune ridges trending northeast to southwest, mobile dunes, and sebkhas, while the coastal fringe has dune calcarenites. The only river to reach the coast is the > Senegal, which has formed a delta at St Louis, where the outflow has been deflected southward by a longshore spit, the Langue de Barbarie. South of Cape Verde there are several large mangrovefringed estuaries in > Guinea Bissau, the > Republic of

Guinea, and > Sierra Leone. Intermittent cliffs are cut in a variety of Shield formations, ranging from Pre-Cambrian granite-gneiss in Liberia and the western Ivory Coast to Palaeozoic sandstones in Ghana (> Liberia, > Ivory Coast, and > Ghana). Basins underlain by Cretaceous and Tertiary rocks form coastal lowlands fringed by sandy barriers, lagoons, and swamps. The coastline includes the broad swampy lobes of the Volta and Niger deltas. The mean spring tide range exceeds 5 m in the estuaries of Guinea-Bissau, but is generally between 1.5 and 2 m. A dominant southerly swell is accompanied by eastward longshore drifting in > Togo and Benin, > Nigeria, > Cameroon, and > Equatorial Guinea. South of the > Congo (> Gabon, Congo, Cabinda and Zaïre and > Angola), narrow coastal plains on basins of Cretaceous and Tertiary rocks are backed by a high hinterland of Pre-Cambrian formations. Beaches are supplied with sand from rivers, notably the Kunene and the Orange, and this west-facing coast receives southwesterly ocean swell that generates northward longshore drifting to build spits (as at Tiger Bay) and supply sand to barriers that enclose lagoons and swamps. The mean spring tide range is generally between 1.5 and 2 m. From southern Angola through > Namibia to the Cape of Good Hope there are sectors of cliff cut in Pre-Cambrian basement and in Cretaceous sedimentary and volcanic rocks, separated by sandy beaches backed by dunes and the Namibian Desert.

Reference Orme AR (2005) Africa, coastal geomorphology. In: Schwartz  ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 13–16

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

13.1  Mauritania

Don Vermeer  

Introduction The Mauritanian coast of West Africa is one of the chief desert coasts of the world. It stretches about 750  km between 16°N and 21°N. Atmospheric subsidence associated with the subtropical high pressure belt accounts for the aridity of those latitudes, but the aridity of coastal Mauritania is intensified by cold-water upwelling and as­­ sociated temperature inversions induced by the Canary Current setting southward at 0.5 knots. Annual precipitation along the coast ranges from less than 50 mm in the north to little more than 300 mm in the vicinity of the Senegal River delta in the south. Three promontories, two within Mauritania and the third to the south in Senegal, anchor this coastal zone and create two distinct sections. The northern section between Cape Blanc and Cape Timiris (>Fig. 13.1.1) is most irregular. Rocky headlands, often more than 20 m in height, reach the sea, and a few large bays, such as Levrier, Arguin, and St. Jean, intrude the coast. Tidal flats margin nearly the whole of this coastal section, the greatest ex­­tent lying north of Cape Timiris and within Uvrier Bay (>Fig. 13.1.2). Mangroves occur only locally, with a northern limit near Cape Timiris (Guilcher and Nicolas 1954). An offshore platform, Arguin Bank, extends 100 km ­seaward, and small portions reach sea level near the outer limits. Arguin Bank shelters the northern section of the coast, greatly reducing wave energy. The coastal zone between Cape Blanc and Cape Timiris is part of the flyway and stopover areas for migratory birds from Greenland, Europe, and tropical Africa, and in 1976 the area was made a national park. The southern section of coastal Mauritania stretches in smooth, nearly unbroken form from Cape Timiris to the Senegal River, and beyond to Cape Verde in Senegal. Neither rocky headlands nor deep embayments mark its gentle arcuate form, and the Senegal River delta protrudes only slightly to deform the coastal outline. The port at Nouakchott consists of a wharf 350 m long, which provides depths of 8 m and permits lighterage. The coast is fronted by a rather steep inner shelf. It is one of the high-energy coasts of the world, and is lined by nearly continuous high dunes.

Western coastal Mauritania lies within the Senegal sedimentary basin that accumulated thousands of me­­ tres of terrestrial and marine sediment during the Creta­ ceous and early Tertiary. Fluvial arkosic sands and gravels (“Continental Terminal”) top the sedimentary sequence, after which a long period of tropical weathering developed an extensive, thick iron-bearing crust (cuirasse ferruginuese). This crust is regarded as the boundary between Tertiary and Quaternary deposits. Pleistocene sea levels during the last interglacial or interstadial reached 4–6 m above the present. Marine regression and extreme aridity followed, and resulted in the formation of linear (seif) red dunes. The trend of the seif dunes varies from about N 20°E in the vicinity of Uvrier Bay to about N 45°E near the Senegal River. With the exception of Cape Arguin, the trend of the dunes declines progressively from true north between Uvrier Bay and the Senegal River. Some of the dunes have blocked the lower Senegal River valley and induced alluvial deposition. Evidence of the Late Quaternary (Holocene) marine transgression in Mauritania comes from mapping of beach materials and the presence of heavy shells (Arca senilis). The Nouakchottian (Flandrian) sea level, about 3 m above present datum, is based on radiocarbon dating of shells, and occurred around 5,500 bp. Stromatolithic algae indicate a marine regression to 3.5 m below present sea level around 4,100 bp, and possibly one or two smaller marine transgressions on this stable coast came between 4,000 and 1,500 bp (Einsele et al. 1974). Raised beaches and beige-and-yellow dunes may be associated with these smaller transgressions, and they often cover much of the Nouakchottian evidence. Furthermore, the yellow sands often lie between the red seif dunes and sufficiently dam the interdune areas to create lakes. The white sands fronting the coast form the present dunes and beaches. Tide ranges are small on the Mauritanian coast. The smallest mean tide range occurs at La Guera on western Cape Blanc (0.8 m) and the greatest at Nouadhibou (1.1 m) within the partially enclosed, shoaling Lévrier Bay. Min­ imal and maximal spring tide ranges occur at La Guera (1.2 m) and Nouadhibou (1.6 m), respectively. Within Uvrier Bay flood and ebb currents are greatest along the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

13.1

Mauritania

⊡⊡Fig. 13.1.2 Southern Mauritanian coast. (Courtesy Geostudies.)

⊡⊡Fig. 13.1.1 Northern Mauritanian coast. (Courtesy Geostudies.) Western Sahara TR EN D N. 20° E.

tidal flat extent 18 meters

Levrie

La Güera 0.8 (1.2)

rocky coast

r Bay

Nouadhibou 1.1 (1.6)

Cape Blanc

21°

barchan dunes

TREND N. 10° E.

linear (seif) dunes 1.0 mean tidal range (1.6) spring tidal range

TR E N. ND 25 ° E .

Cape Arguin 1.1 Arguin Bay (1.4)

Mauritania

Arguin Bank

Cape Tafarit

20°

Cape Timiris

TR

Ocean

N.

St. Jean Bay

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EN 40 D ° E.

Tidra

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Tenioubrar Sebkra

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0 km 17°

16°

western side. At Nouadhibou the flood current sets northward with a maximum velocity of 1.5 knots, and the ebb current sets southward with a maximum velocity of 2.5 knots. Shoals within Uvrier Bay are more extensive on its eastern side than on its western side. Mauritania has active dunes, which drifts from land to sea, supply­­ing sand to an open ocean coast and continental shelf, and hence contributing to beaches and marine sediment (Sarnthein and Diester-Haass 1977). Barchan dunes reach the coast between northern Levrier Bay and Cape Tafarit (>Fig. 13.1.1); they seldom exist at the coast south of Cape Tafarit. Northerly and northeasterly trade winds dominate this coastal zone, and account for the alignment of the barchan and linear dunes. This belt currently contributes between 5 and 13 million m3 of quartzose sand to

the continental shelf each year (Sarnthein and Walger 1974). The relatively coarse sand is trapped and piled up as sand wedges that laterally prograde from the shore with the upper edge at sea level. The wedge gradually widens the foreshore. During the Pleistocene lowering of sea level, the sand wedges would have formed near the upper part of the continental slope and would have moved to the ocean basins by turbidity currents or sand flows, or both. During

Mauritania

13.1

⊡⊡Fig. 13.1.3 Coastal dunes about 20 m high, about 40 km north of Nouakchott. Wreck of wooden vessel embedded in beach in foreground. (Courtesy J.R. Illick.)

Pleistocene peak glacial periods, when the linear (self) dunes were active and not fixed as they are today, sand discharge to the Atlantic Ocean is estimated to have been at least five to ten times the present rate. The Pleistocene sand contributions were comparable to the present total loads of rivers such as the Nile or Niger. Sand from active barchan dunes and erosion of fossil seif dunes provides material that is re­­worked and moved progressively southward by longshore drift through the beach and dune systems of the coast. Northwesterly to northerly waves strike the coastal zone more than 65% of the time, and especially stormy conditions prevail during the winter months as the result of cyclonic disturbances that pass to the north. Sand drifting southward contributes to coastal progradation in southern Mauritania, and moves around the head of steep submarine canyons to reach as far south as Cape Verde in Senegal (Sarnthein and Diester Haass 1977). The steep inner shelf and exposure to the Atlantic swell permit wave energy of enormous magnitude (1,124 million ergs/s) on the coast; more wave energy arrives on the southern coastal area in three hours than hits the Mississippi delta in 365 days (Wright and Coleman 1973). The high wave energy and southward longshore drift have flattened the Senegal River delta and turned the river southward (>Fig. 13.1.2). The river flows parallel with the coast behind an elongated sandy spit on which St. Louis is built, before it discharges its load into the ocean.

Coastal dunes are poorly developed north of Cape Timiris, but well developed to the south. Immediately south of Cape Timiris, coastal dunes attain heights of 35 m or more, while farther south their heights are generally less than 20 m (>Fig.  13.1.3). The width of the coastal dune belt varies from 200 to 500 m but reaches a maximum of 2 km north of the Senegal River delta. Where coastal dunes truncate seif dunes south of Cape Timiris, salt pans (sebkhas, elsewhere known as sabkhas) occupy the interdune areas. Farther south, the sebkhas paral­­lel the coast and lie immediately behind the dunes. A few are large, such as Ndaghamcha Sebkha, while others north of the Senegal River delta flood during peak river discharge, usually in September or October. They continue south of the delta into Senegal, where they are known as niayes.

References Einsele G, Herm D, Schwarz HU (1974) Holocene sea level fluctuations at  the coast of Mauritania. Meteor Forschungsergebnisse, Reihe C 18:43–62 Sarnthein M, Diester-Haass L (1977) Eolian-sand turbidites. J Sediment Petrol 47:868–890 Sarnthein M, Walger E (1974) Der äolische Sandstrom aus der W Sahara zur Atlantikküste. Geolo Rundsch 63:1065–1087 Wright LD, Coleman JM (1973) Variations in morphology of major river deltas as functions of ocean wave and river discharge regime. Bull Am Assoc Petrol Geol 57:370–398

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13.2 Senegal and Gambia André Guilcher

1.  Introduction

enclave within Senegal, includes only the immediate surroundings of the lower course of the Gambia River (>Fig.  13.2.1). The province of Senegal situated to the south of Gambia is called Casamance from the name of

Senegal and Gambia are considered together in this de­­ scription of their coasts, because Gambia, a small country

⊡⊡ Fig. 13.2.1 The coast of Senegal and Gambia.

z ki R ke a L

Marigot des Maringouins

N

ke



ne

gal

R iv er

ie

r

La

D E L T A

Se

a

r Ca ny

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Cô te Joal

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DAKAR Cap de Naze

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Rufisque

tit

Ve rt

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Cap

i

Kaya

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Saint Louis

Saloum Kaolack

River

Pointe de Sangomar BANJUL Bald

Cape

Solifor

Pointe aux Oiseaux

Gam

bi

a Riv er

n Ca s ama ce Ri ve

r

Cape

Ziguinchor

0

100 km

Cap Skirring Sandy coast State boundary

Cliff

tidal marsh (not shown in Senegal Delta)

limit of Nouakchottian (Holocene) transgressior

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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13.2    

Senegal and Gambia

⊡⊡ Fig. 13.2.2 Small Rhizophora and herbaceous marine vegetation on the Senegal delta near Saint Louis.

the river that flows through it. Senegal has a 531 km coastline, The Gambia, 80 km. The geological structure of these countries consists of  a Mesozoic and Cainozoic basin on the western edge of  the Pre-cambrian and Palaeozoic African block. The sedimentary series begins with the Upper Jurassic, and includes mostly Cretaceous sands, sandstones, clays, and marls, Eocene limestones, clays, and phosphates, and Tertiary sandstones belonging to the so-called Terminal Con­tinental, the precise age of which is not well known. The main coastal outcrops of these rocks (especially the Eocene) are between Dakar and Sangomar Point. At the westernmost tip of the heavily faulted Cape Verde peninsula, volcanic rocks of Tertiary and Quaternary age are found. About 75,000 years ago, during the Late Pleistocene, the so-called Inchirian marine transgression left coastal deposits 20 m beneath the mouth of the Senegal River. A succeeding regression took the sea down to about 120 m below present sea level about 18,000 years ago, in an arid phase when the Ogolian red dune belt spread over a large part of western Senegal. The Late Quaternary (Flandrian) marine transgression, here known as the Nouakchot­­ tian, reached its maximum 5,500 years ago when the sea invaded the Senegal, Sine, Saloum, Gambia, and Casamance valleys. The gulfs thus created have since been largely filled with fluvial sediment, together with some sediment swept in from the sea. Details of a dated succession of sea level changes were given by Giresse et al. (2000). The modern Senegal delta buried the Nouakchottian marine deposits in the north, where successive coastal

spits enclosed lagoons and prograded the coastline. Some of the highest Nouakchottian beaches, however, are still exposed to the rocky coasts of the Cape Verde peninsula and in several places along the Gambian coast (Michel 1973). The climate, which governs the river discharge, is of tropical boreal type, with a long dry season in winter and rains concentrated in summer. The rainfall and the length of the wet season increase from north to south: in the 1931–1960 period mean annual rainfall was 370 mm at Saint Louis, 670 mm at Dakar, 800 mm at Kaolack, and 1550 mm at Ziguinchor. In recent years, the country has experienced severe droughts, so that the mean annual rainfall for 1970–1980 dropped to 208 mm at Saint Louis, 345 mm at Dakar, and 1270 mm at Oussouye near Ziguinchor. The annual discharge from the Senegal and Gambia rivers shows a prominent peak in September and very low levels from January to July. The Senegal discharge during floods has varied from 2500 to 3700 m3/s, and the mean ratio between the extreme months is 1:860. For this reason, and because of the very low gradient of this river, tidal oscillations are observed as far as 440 km upstream during the low flow stages, the salinity rising to 29 parts per thousand 60 km upstream. Mean spring tide ranges are small: 1.2 m at Saint Louis, 1.3 m at Dakar, and 1.6 m at Banjul (formerly Bathurst) in Gambia. In the Guier and Rkiz lakes, which receive part of the annual river floods, water level fluctuates accordingly; during the larger floods, the Senegal spreads into coastal lagoons such as Marigot des Maringouins, which extends

Senegal and Gambia  

⊡⊡ Fig. 13.2.3 The lower course of the Senegal River and the Langue de Barbarie.

   13.2

reaching the whole coast of Senegal and Gambia as far south as Casamance. The southwesterly swell is extremely rare (less than 2% of observations). Because of the outline of the coast, the northwest swell is more effective north of Cape Verde than to the south. The dominant onshore wind, which is moderate, also blows from the NNW, and produces waves that accompany the northwest swell, causing southward flow of longshore currents and southward longshore drifting of beach material.

2. The Senegal Delta and the Langue de Barbarie

behind the Mauritanian coastline. Much less information is available for the Gambia River. Wave action is dominated by the northwest swell, ­generated in the storm belt of temperate latitudes, and

The distributaries of the Senegal delta are bordered by postNouakchottian natural levees. In smaller creeks, the only currents in the dry season are the ebbing and flowing tidal currents. At this latitude (16°N) on the African west coast, the swamp vegetation (>Fig.  13.2.2) consists principally of  herbaceous plants, but mangroves (Avicennia and Rhizophora) also grow patchily along the banks of the river upstream to Saint Louis; the trees are rather stunted because they are close to their northernmost limit (near Cape Timiris in Mauritania) (Guilcher and Nicolas 1954). On the ocean coast, beach sand drifts southward under the influence of the northwest swell, arriving obliquely to the shore. A pattern of small oblique spits and bars and intervening troughs is typical of this beach, and as on the coast of the Landes of Gascoyne in France, it is an adaptation of the nearshore topography to the oblique swell, with rip currents occurring in the troughs. The amount of sand in transit in front of Saint Louis has been estimated at 500,000–900,000 tonnes per year. The result is that the Senegal mouth has been deflected far to the south by a longshore sand spit called the Barbary Tongue (Langue de Barbarie). The length of this spit fluctuates over time, but the outlet from the Senegal River can be up to 25 km south of Saint Louis (>Fig.  13.2.3), while the northernmost known location of the river mouth was 2.5 km north of Saint Louis in 1850. Twenty-four breaches of the Langue de Barbarie were reported between 1850 and 1980, resulting either from heavy swell on the ocean coast or scour from the river on the inner shore during floods. After each episode of truncation, the Langue has grown again to the south. There has also been slight retreat of this barrier coastline (Barusseau et al. 1996). The moderate local NNW winds affect the coastal area  in two ways. First, they build low dunes oblique to the  coastline, periodically destroyed on the Langue de Barbarie, but more stable elsewhere. Second, they generate small waves across the lower Senegal River and its parallel distributaries, and these build sets of small evenly spaced

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924

13.2    

Senegal and Gambia

⊡⊡ Fig. 13.2.4 Tidal creeks in Casamance.

recurved southward spits of the Azovian type (as defined on the shores of the > Sea of Azov), particularly in the vicinity of Babagueye Island. From the Senegal mouth to Cape Verde is the so-called Niayes coast, a sandy coast with low dunes, behind which is a series of depressions, the niayes, which are swamps or lakes during the wet season and dry pans in winter. Much of the sand that drifts southward along this coast probably flows out into a submarine canyon, the Kayar Canyon, the head of which lies very close to the coast. To the south is Cape Verde Peninsula, a bold promontory west of Dakar, where there are cliffs exposing dolerites and basalts. At Ouakam, the dolerites show ­corrosion forms with a seaward rampart at the foot of a

cliff 25 m high (Guilcher et al. 1962). There are small islands and stacks of columnar basalt. Northwest swell breaks heavily on the northern coast of Cape Verde, and are refracted round it toward the Petite Côte, between Rufisque and Joal. There are cliffs cut in clays, limestones, and clayey sandstones, and although the coast is comparatively sheltered from the northwest, they are nevertheless retreating. At Cap de Naze, where the cliff is exceptionally high (55 m), erosion is accompanied by slumping and gullying. In the shelter of Cape Verde, the southward beach drifting diminishes to only 30,000–70,000 m3/year. It increases southward to about 300,000 m3/year, and has supplied a spit, Pointe de Sangomar, at the southern end of

Senegal and Gambia  

the Petite Côte. The growth of this irregular sandy spit amounted to some 3,300 m from 1895 to 1980, with a decrease since 1930. South of this spit are several large mangrove-fringed estuaries, the Southern Rivers: Sine-Saloum, the Gambia River, and Casamance River. A broad bay with mangrove fringes and deltaic islands extends past the Sine-Saloum estuary, the Holocene evolution of which has been studied by Ausseil-Badie et al. (1991). Banjul is on the southern side of the Gambia River, and beyond this are cliffs and rocky platforms cut into sandstones of the Terminal Continental, forming Cape Bald and Cape Solifor in Gambia. The Gambia River does not have swampy shores and islands of the kind seen on the Sine-Saloum estuary to the north, because the Gambia estuary is incised into sandstones capped by a ferruginous crust, and hence is generally narrower (Michel 1973). The muddy swamps of the Sine-Saloum and Casa­ mance Rivers, fed with seawater by intricate sets of creeks, bear many more thriving mangroves than the Senegal delta because of the more southern position and higher rainfall. However, in these green intracoastal landscapes, there are muddy areas devoid of vegetation, called tannes, between the mangroves and the terrestrial vegetation inland. During the dry season, sea water covers all the tidal flats at spring tides, but at neap tides the inner parts are not submerged for several days; because there is no rain, ­salinity increases in these areas by evaporation and becomes too high for mangrove growth. On the other hand, these tidal marshes have been partly reclaimed and transformed in polders for

   13.2

rice cultivation by the Diola population over a very long period, so that the landscape is partly artificial. At the southern border of Gambia, the cliffed coast gives place to a set-back sandy beach, which runs south to another spit, the Point aux Oiseaux, which is growing slowly southward. Beyond it are more mangrove swamps, islands, and intersecting channels at the broad mouth of Casamance River (>Fig.  13.2.4). Further south are more cliffs in Tertiary sandstones and clays, the coast swinging ­south-east at Cape Skirring, toward the border of Guinea Bissau.

References Ausseil-Badie J, Giresse P, Pazdur M (1991) Holocene deltaic sequence in the Sakoum estuary, Senegal. Quat Res 36:178–194 Barusseau JP, Descamps C, Golf A (1996) Evidence for short term retreat of the barrier shoreline. Quat Sci Rev 15:763–771 Giresse P, Barusseau JP, Causse C, Dioul B (2000) Successions of sea-level changes during the Pleistocene in Mauritania and Senegal distinguished by sedimentary facies study and U/Th dating. Mar Geol 170:123–139 Guilcher A, Nicolas JP (1954) Observations on the Langue de Barbarie and the Senegal River near Saint Louis (in French). Bulletin d’Information, Comité Central d’Océanographique, (Paris) 6:227–242 Guilcher A, Berthois L, Battistini R (1962) Shore corrosion forms on volcanic rocks, particularly in Madagascar and at Cape Verde (Senegal). CahOcéanograph 14:208–240 Michel P (1973) A geomorphological study of the basins of the Senegal and Gambier Rivers (in French). Office de Recherche Scientifique et Technique d’Outre-mer Mémoir 63, Paris

925

13.3 Guinea Bissau

E.S. Diop

Introduction The coastline of the Republic of Guinea Bissau is about 350 km long, and consists of an irregular mainland with numerous rias and the estuaries of the Cacheu and Geba rivers, and a group of offshore islands, the Bissagos (Bijagos) Archipelago. The coastal plain is dissected by an intricate network of mangrove-fringed tidal channels, and the outer shores are sandy, with many shoals (Berthois 1958). Wave energy is generally low, partly because the northwest and southwest Atlantic swells are dissipated across a broad sector of continental shelf, and partly because of the protection of the archipelago. Tides are semidiurnal. They are augmented across the broad continental shelf and into the funnel-shaped estuaries and inlets, so that Cacheu in northern Guinea Bissau has a mean spring tide range of 2.8 m, while Bissau, on the north coast of the Geba estuary, has a mean spring tide range of 5.1 m, increasing to 6.4 m at Port Gole. Bubaque, in the archipelago, has a mean spring tide range of 4.6 m. Guinea Bissau has a coastal climate that includes a dry season (between November and May) when northwest swells predominate and the prevailing winds are north– northwesterlies (trade winds), and a rainy season (between June and October) when southwest swells prevail and southwest winds (monsoon winds) are dominant. These seasonal alternations influence the sedimentary pattern of the coastal and shelf areas. In the dry season progradation occurs, the northwest swells producing longshore drifting to the south, prolonging sand bars, and moving shoals southward across estuaries and inlets; in the wet season erosion prevails. A lateritic (ferruginous) duricrust is extensive in the hinterland, and comes to the coast locally to form low cliffs, as near Caio on Jeta Island and on parts of the Bissagos Archipelago. Elsewhere, the duricrust descends to outcrop in the intertidal zone, where mangroves (Rhizophora and Avicenna) grow sparsely, and continues to a depth of several metres on the inner continental shelf. Plio-Quaternary formations outcrop on these ria coasts (Ayme 1965), which were formed when the Late

Quaternary (Flandrian or Nouakchottian) marine transgression submerged an undulating coastal lowland with  many valleys. This transgression came to an end about 6,000 years ago, and there was a brief phase when  the sea rose about 2 m above its present level. Holocene deposition of sandy, muddy, and organic sediment has blanketed the early and middle Quater­nary,  as  well as older formations. The Holocene formations ­continue on to the continental shelf, where there are ­submerged deltas, submarine shoals, and former river channels, and deposition is still proceeding, with large quantities of dark grey silt and clay being delivered by the rivers. Coastal waters are turbid, with much sediment is suspension, particularly after rainy periods when the rivers discharge plumes of sediment-laden water. The Guinea Bissau coast has the following geomorphological units. There are mangrove swamps and mud­­ flats with salt marshes, including herbaceous tannes, a kind of terrace formed on old mudflats. There are sandy beaches, sometimes backed by beach ridges and low dunes with intervening swales, which may be swampy. There are nearshore sand bars and muddy shoals, extensively exposed at low tide, and there are a few low cliffs and grassy bluffs. The estuaries and tidal inlets include meandering and intersecting adjacent channels, with some arcuate lagoons that formed as cut-off meanders. Mangrove swamps and mudflats predominate. They have formed as the result of the deposition of fine-grained muddy sediment (blue and grey silt and clay, often associated with remains of marine organisms). The accreting mudflats have been colonised by mangroves, which are adapted to grow in sea water subject to frequent tidal oscillations. Most of the mangroves are of three species: Rhizophora racemosa, Rhizophora mangle, and Avicennia africana. The mangroves are backed by tannes, a kind of sandy mudflat slightly above normal high tide level, which may be bare and saline, sometimes with white salt encrustations, or (above normal high tide level) covered with herbaceous vegetation. In the more saline areas this includes the halophytes Philoxerus vermicularis, Sesuvium portulacastrum, and Paspalum vaginatum; in the less saline

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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13.3

Guinea Bissau

environments Heleocharis plantiginea and Scleria racemosa are found. On the west and southwest coasts of the archipelago, where wave action is relatively strong, cliffs and bluffs locally interrupt the mangrove fringe and rock outcrops appear in the form of shore reefs and platforms. Sandy beaches sometimes develop on longshore bars built by northwest or southwest swells, and generally extend in a straight line and are parallel to the coast. They form northwest of the Cacheu River and south of Bolama. They contain heavy minerals such as rutile, zircon, and ilmenite, which are also found in the other Quaternary deposits. In general, when these sandy beaches are built on longshore bars, they may develop mobile sand dunes. Some sandy beaches are backed by older parallel beach ridges with intervening swales that run north–northwest to south–southeast in narrow, discontinuous series. They result from coastal progradation, mainly during the

emergence when the sea fell back from the slightly higher mid-Holocene stand. They are sometimes capped by dunes, with crests rising to 7.5 m. Sandy and muddy shoals are found in the estuaries and tidal inlets (Guilcher 1979), where they are exposed at low tide. They are accreting and unstable, with intervening channels that migrate laterally, and they frequently obstruct navigation.

References Ayme JM (1965) The Senegal salt basin. In: Salt Basins around Africa. Petroleum Institute, Amsterdam, Elsevier, London, pp 83–90 Berthois L (1958) Formation of estuaries and deltas (in French). Academe Sci Comptes Rendus 247:947–950 Guilcher A (1979) Marshes and estuaries in different latitudes. Interdiscipl Sci Rev 4:158–168

13.4 Republic of Guinea

E.S. Diop

Introduction The coastline of the Republic of Guinea is about 300 km long. Much less intricate than the coast of Guinea Bissau, it is mainly low-lying and dominated by rias, tidal inlets, and small estuaries (Berthois 1958). Mangroves are extensive around these, but the open coast is largely beachfringed. Plio-Quaternary formations outcrop in the coastal region, the underlying Lower Palaeozoic rocks rising in the hinterland and in the ridge that runs out to Cape Verga, a cliffed promontory. There are also cliffs bordering the basic and ultrabasic intrusive rocks of the Kaloum Peninsula and the outlying Los Islands. Ferruginous duricrusts cap cliffs and outcrop as shore platforms, as at Conackry, where scattered mangroves grow on them. The rias, tidal inlets, and estuaries were formed by ­partial submergence of an undulating Pleistocene lowland  incised by river valleys during the late Quaternary (Flandrian or Nouakchottian) marine transgression. This attained the present sea level in Holocene times, about 6,000 years ago, and then rose briefly, about 2 m above this level, before falling back in a phase of mid-Holocene emergence. There has subsequently been extensive deposition of sand and mud on the coastal plain, and on the continental shelf, which is here relatively wide (about 140 km). There are submerged deltas, submarine shoals, and former river channels offshore, the Konkoure River having a submerged channel that continues across the continental shelf (MacMaster et al. 1970). There are also submarine deltas off the mouths of the Kapatchez and Pango rivers. Wave energy is generally low because the southwest swell from the Atlantic is dissipated across the broad continental shelf. Tides are semidiurnal, and augmented as they cross the wide continental shelf: Conackry has a mean spring tide range of 3.2 m. The Republic of Guinea has a tropical coastal climate that includes a dry season (between December and April) when the prevailing winds are north–northwest (trade winds), blowing off the land, and a rainy season (between May and November) when southwest swells prevail and  southwest winds (monsoons) are dominant. These

s­ easonal alternations influence coastal processes, with erosion in the wet season and accretion in the dry season. Coastal waters are turbid, with much sediment in suspension, particularly after rainy periods when the rivers discharge plumes of sediment-laden water. Surface turbidity in the offshore Conakry is 10–20 mg/L in the dry season, increasing to several dg/L in the rainy season; the bottom turbidities are four or five times greater. These turbidity values are very high, and it is not surprising that siltation in inlets and estuaries impedes navigation. The harbour on the Kapatchez River was abandoned because of siltation. The Republic of Guinea coast has the following geomorphological units. There are mangrove swamps and mudflats with salt marshes, including herbaceous tannes, a kind of saline terrace formed on old mudflats, like those in Guinea Bissau, but on a smaller scale. There are sandy beaches, sometimes backed by beach ridges and low dunes with intervening swales, which may be swampy. There are nearshore sand bars and muddy shoals, extensively exposed at low tide, and there are a few low cliffs and grassy bluffs. Mangrove swamps and mudflats are extensive alongside inlets and estuaries. They have formed as the result of the deposition of blue and grey silt and clay carried down rivers, often with remains of marine organisms. The accreting mudflats have been colonised by mangroves, which are adapted to grow in seawater subject to frequent tidal oscillations. Most of the mangroves are of three species: Rhizophora racemosa, Rhizophora mangle and Avicennia nitida (africana). The mangroves are backed by tannes, a kind of sandy mudflat slightly above normal high tide level, which may be bare and saline, sometimes with white salt encrustations, or (above normal high tide level) covered with herbaceous vegetation. In the more saline areas this includes the halophytes Philoxerus vermicularis, Sesuvium portulacastrum, and Paspalum vaginatum, while in the less saline environments Heleocharis plantiginea and Scleria racemosa are found. Sandy beaches are extensive on the coast of the Republic of Guinea, shaped by southwest swell that has been weakened in crossing the wide continental shelf.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Some sandy beaches are backed by older parallel beach ridges with intervening swales that run north–northwest to south–southeast. They result from coastal progradation, mainly during the emergence when the sea fell back from the slightly higher mid-Holocene stand. They are sometimes capped by dunes, with crests rising to about 6 m. Sandy and muddy shoals are found in the estuaries and tidal inlets (Guilcher 1979), where they are exposed at low tide. They are accreting and variable, with intervening channels that migrate laterally, and they frequently obstruct navigation, as in the roadstead (channel) of Conakry, where dredging of 750,000 m3/year is necessary to maintain the port.

Bluffs and cliffs are localised around Verga Cape and the Kaloum Peninsula, where they are exposed to long swells and wave abrasion. There are gravelly shore platforms in front of them.

References Berthois L (1958) The formation of estuaries and deltas. Academie de Science, Comptes Rendus 247:947–950 Guilcher A (1979) Marshes and estuaries in different latitudes. Interdiscipl Sci Rev 4:158–168 MacMaster RL, Lachance TP, Ashraf A (1970) Continental shelf geomorphic features of Portuguese Guinea, Guinea and Sierra Leone (West Africa). Mar Geol 9:203–213

13.5 Sierra Leone

1.  Introduction The coastline of Sierra Leone is 402 km long (>Fig. 13.5.1). The western coast is mainly low-lying and dominated by mangrove-fringed rias, apart from the low cliffs of the Bullom Peninsula and the mountainous Freetown Peninsula. The climate is humid tropical, with a wet season (southwest monsoon) during May–November and a dry season during December–April. Freetown has a mean temperature of 26.7°C in January and 25.6°C in July, with a mean annual rainfall of 3,434 mm. Wave energy is relatively low, with an occasional much-refracted weak northwesterly swell from the North Atlantic in the dry season, and mainly low waves of no fixed direction. Locally generated storm waves occur during the wet season. The eastern coast between the Turtle Islands and the Liberian border is dominated by the southwesterly Atlantic swell. The tides are semidiurnal. The mean spring tide range at Freetown harbour is 2.8 m, and is slightly higher on the eastern coast. The rias on the western coast were formed by partial submergence of an undulating Pleistocene lowland incised by river valleys during the late Quaternary (Flandrian) marine transgression. This attained present sea level in the Holocene, about 6,000 years ago, and then rose briefly about 2 m above this level, before falling back in a phase of mid-Holocene emergence. There has subsequently been extensive deposition of sand and mud on the coast and the continental shelf, which is relatively wide here (about 100 km). Coastal waters are turbid, with much sediment in ­suspension, particularly after rainy periods when the ­rivers discharge plumes of sediment-laden water. Sedi­­ ment in the estuaries consist of sand bars and mudflats, with gravel in the channels (Tucker 1973). Mangrove swamps and mudflats are extensive alongside rias and estuaries. They have formed as the result of the deposition of silt and clay carried down rivers. The accreting mudflats have been colonised by mangroves, notably Rhizophora species.

The eastern coast is largely beach-fringed, with numerous parallel beach ridges on the barriers of Sherbro Island and the Turner Peninsula.

2.  The Sierra Leone Coastline South of the Republic of Guinea border, extensive mangrove swamps border the ria fed by the Great and Little Scarcies rivers. The mangroves belong to the four genera of the occidental mangrove province. The most common, Rhizophora (a stilt root mangrove), is represented by three species. All are found in the part of the intertidal zone that is flooded every day, but R. racemosa prefers the siltier substratum near the low water mark and colonises newly emerged mudflats, while R. harrisonii and R. mangle are found on slightly higher ground. Nowhere does Rhizophora reach its maximum possible height, and it does not thrive in sandy substratum, which is frequently occupied by Avicennia nitida. This species is usually found in stands higher up the intertidal zone, where it is not generally submerged every day; it is more salt-tolerant than Rhizophora, and rarely grows higher than 5 m. Around the high spring tide line are found the two other genera, each represented by one species, Conocarpus erectus and Laguncularia racemosa, usually growing as scattered individuals in a thicket. Between the Great and Little Scarcies rivers, much of the area of the former mangrove swamp has been embanked to form polders that are used for rice cultivation. To the south the Bullom Peninsula has a small area of Holocene beach ridges built by waves, but modified by current action. Beyond these are cliffs up to 20 m high, cut into Eocene marine and estuarine deposits. These consist of a series of poorly consolidated, near-horizontal gravels and clays, partly laterised in places. The cliffs cut are almost vertical because erosion has been rapid: falls and slides are common, especially in the wet season, and the fallen debris is removed from the cliff foot by wave action. This fallen

This is a revised version of a chapter by P.A. Scott in The World’s Coastline (1985: 569–573).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 13.5.1 Coastal landforms of Sierra Leone.

Sierra Leone  

material becomes partly indurated on contact with seawater. A narrow beach of sand and scattered shingle is found at the base of the cliffs. More mangrove swamps fringe the ria into which the  Rokel River flows. Evidence of drowned river channels  points to a Late Quaternary marine transgression in  the Rokel estuary. To the south is the mountainous Freetown Peninsula, rising to 930 m. This is the eastern margin of a much eroded and submerged Triassic intrusion of gabbro-norite rocks, which are generally resistant to both mechanical and chemical erosion in coastal ­outcrops. The coastline of the peninsula is a succession of headlands separating beach compartments. The headlands are steep, thickly vegetated slopes down to the spray zone just above high tide level, and shore platforms are poorly developed. Many of the longer beaches show longshore drifting of sand, some in a northerly direction, as at Tokeh, and others to the south. Spits divert the mouths of many small mountain streams, and enclose mangrove-fringed lagoons. The peninsula beaches consists of medium to fine  quartz sand, which has evidently come from the sea  floor during and since the Late Quaternary marine transgression. The profiles of the peninsula beaches vary seasonally. During the dry season they show well-developed berm forms, bordered seaward by large beach cusps, spaced at 25–35 m (Worrell 1969). In the wet season the berm is eroded, the beach flattened, and a low cliff is cut in sand at the high water mark. This is usually buried during the following dry season. This seasonal alternation of erosion and deposition appears to be balanced, except where beach sand has been extracted for building purposes, where erosion rates are high and there is no dry season build-up. Behind many of the peninsula beaches and lagoons, and extending up the larger river valleys, is a gently sloping abrasion platform rising to a marked break of slope, 50 m above sea level and backed by mountain slopes. Its seaward edge is usually a low cliff, either at the sea or behind a lagoon. Gregory (1962) called these platforms raised beaches, and although their surfaces do not bear any beach materials (they are laterised in most cases), they are certainly platforms of marine origin. At the southern end of the Freetown Peninsula the Banana Islands run offshore, also composed of Triassic gabbro-norite. The bay to the south is mangrove-fringed, round to the Shenge Peninsula, which has low cliffs cut into the Eocene Bullom Sediment that recur here. The features are similar to those on the Bullom Peninsula, but the cliffs in horizontally bedded Eocene clays are lower here (>Fig. 13.5.2).

   13.5

Mangroves border rias that extend behind the Shenge Peninsula, and there is a small area of beach ridges on the southern coast. The Sherbro River has a broad outlet in the bay to the south, bordered by extensive mangroves. Rhizophora racemosa grows up to 20 m high in favourable locations here. On the southern side is Sherbro Island (Anthony 1991, 1996), a low-lying barrier island with numerous parallel beach ridges, backed by mangroves and bordered seaward by a wide beach washed by southerly ocean swell (Worrell 1970). The beach ridges have been truncated on the western shore: they have no recurved terminations, and the Turtle Islands may be relics of the truncated barrier. These are mostly vegetated sand islands, the smaller ones constantly shifting and changing in size, and occasionally disappearing altogether. Barrier truncation may have resulted from tectonic subsidence in the area south of the Shenge Peninsula, for mangroves have invaded the swales between beach ridges in the western part of Sherbro Island. Interrupted by a southern outlet from Sherbro River, the Sherbro Island barrier continues eastward as the Turner Peninsula. The beach ridges are built of quartz sand supplied by the rivers of Liberia and southern Sierra Leone; the sand has drifted eastward. Two sets of ridges occur: a narrow series of closely-spaced Holocene ridges still forming along the ocean coastline, and, behind them, a much more widely-spaced series thought to be of Pleistocene age, although no dates are available. The Holocene ridges are about 70 m wide, with swales between of about the same width, which flood with fresh water in the wet season, preventing tree growth. Palms and low shrubs grow on the ridges. The much more widely spaced older ridges have been dissected by streams, and show limited soil development. The Sewa and Jong rivers and their tributaries are diverted westward by the beach ridges, the lower course of the Sewa cutting through the older ridges to enter a series of freshwater lakes behind the Holocene beach ridges. The beaches on the Atlantic coast of Sherbro Island and the Turner Peninsula are wide, and appear to be ­prograding (>Fig.  13.5.3). The beach sand is coarse to medium in texture, with a mean diametre of 0.5 mm. The beaches are wider and steeper than those of the Freetown Peninsula. The sandy beach and beach ridges run eastward to a mangrove-fringed inlet at the mouth of Moa River, which has a complex of sand bars at the entrance, and on to the Mano estuary, the mouth of which has been deflected westward by a spit. The Liberian border runs along the Mano River.

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⊡⊡ Fig. 13.5.2 Receding cliffs at Shenge, cut in almost ­horizontally bedded clays. Although the photograph was taken at the end of the dry season, there is no sign of berm accretion.

⊡⊡ Fig. 13.5.3 The wide sandy beach on the Atlantic coast of Sherbro Island is accreting (contrast > Fig. 13.5.2).

References Anthony EJ (1991) Beach-ridge plain development: Sherbro Island, Sierra Leone. Z Geomorphol, Suppl.Bd 81:85–98 Anthony EJ (1996) Evolution of estuarine shoreline systems in Sierra Leone. In: Nordstrom K, Roman I (eds) Estuarine shores: evolution, environments and human alterations. Wiley, Chichester, pp 39–61

Gregory S (1962) The raised beaches of the peninsular area of Sierra Leone. T Brit Geogr 31:15–22 Tucker ME (1973) The Sediment of tropical African estuaries, Freetown Peninsular, Sierra Leone. Geol Mijnbouw 52:203–215 Worrell GA (1969) Present day and subfossil beach cusps on the West African coast. J Geol 77:484–487 Worrell GA (1970) Sherbro Island, Sierra Leone. Bulletin de l’Institut France d’Afrique Noire, Series A 32:1–17

13.6 Liberia

1.  Introduction Stretching about 579 km and facing SW, the Liberian coast (long known to mariners as the Grain Coast because of trade in Malagueta pepper) has a mélange of rocky outcrops, some of which form conspicuous headlands, interspersed with long sandy beaches, some of which have become barriers backed by tidal lagoons. There are few good roadstead anchorages and almost no natural harbours. Liberia has a humid tropical climate with year-round rainfall rising to a maximum in the wet season (May– October). Monrovia has mean monthly temperatures of 26.1°C in January and 24.4°C in July and a mean annual rainfall of 5,138 mm. Tide ranges are small: Monrovia has a mean spring tide range of 1.1 m. The continental shelf is narrow, and the SW swell that originates in the South Atlantic breaks heavily on the shore, forming sandy beaches on which giant beach cusps are often seen. Longshore drifting is generally to the southeast, but reverses when southerly winds produce waves that arrive obliquely to the shore from that direction. The coastal zone is up to 50 km wide, and has an extensive cover of unconsolidated sand, clay, and gravel deposits of Quaternary age (Schulze 1973). Often found in proximity to rivers flowing perpendicular to the coast, these deposits also form some of the emerged beaches found especially in the Monrovia area. The Quaternary deposits partially cover intrusive diabase of Palaeozoic age, and there are scattered outcrops of Tertiary sandstone between the entrances to the Mesurado and St. John Rivers and southeast of the city of Monrovia (White 1972). The more resistant diabase dikes form conspicuous promontories and cliffs. Weathered remnants of those dikes and earlier laccoliths form numerous rocky outcrops and offshore rocks, and as weathering appears to be mainly by solution their appearance is one of rain-wash flutings or lapies. Details of features produced by weathering processes on rocky shore outcrops of dolerite at Mamba Point were described by Tricart (1962).

A powerful surf and exceptionally strong rip currents characterise most of the coastline, except where offshore rocks are sufficiently dense to disrupt surf and currents. There are few islands.

2.  The Liberian Coastline From the Sierra Leone border at the mouth of Mano River, a beach-fringed coast curves south to Cape Mount (305 m), the highest point on the Liberian coast. Behind this mountain is Lake Piso, a brackish water tidal lagoon generally less than 5 m deep. Cape Mount is linked to the mainland by a depositional isthmus (tombolo), which enclosed the former bay that became Lake Piso. To the south east, a sandy beach extends to Cape Mesurado. The Lofa River enters the sea by crossing a zone of coastal lagoons to a gap in the sandy barrier, and St Paul River, navigable to White Plains, some 32 km inland, is generally blocked by sand bars and shoals deposited by swell at the entrance, but there is an outlet though Stockton Creek to the mouth of the Mesurado River. The coast along this stretch is fringed with rocks and rocky points interspersed with stretches of light brown sand. Cape Mesurado (91 m) at Monrovia has been quarried to obtain construction material for two 1.6 km jetties built between 1944 and 1947 to enclose the port of Monrovia. The coastal cliffs are partly quarry faces. Southeast of Cape Mesurado is a large tidal lagoon through which the Du and Farmington rivers flow to the sea, and 2 km to southeast of the entrance is a headland of diabase. Further along the coast is the similar St. John River lagoon system (>Fig. 13.6.1). The port of Buchanan, SE of the St. John River, is formed by jetties 1920 m and 610 m long, built of rock quarried from a nearby diabase dike. As at Monrovia, the port site was chosen because of the availability of nearby construction material. It has a major iron ore shipping facility. The washing of iron ore near the port has caused massive discoloration of the ocean in a zone several hundred metres wide along the coast from the St. John River

Edited version of a chapter by W. R. Stanley in The World’s Coastline (1985: 575–583). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 13.6.1 Lagoons and barrier spits at the mouth of St John River. Ocean swell has shaped the spit, and the entrance is maintained by river outflow, which disrupts the wave pattern.

⊡⊡ Fig. 13.6.2 Barrier and lagoon near Aruan Point, SE of Greenville.

to Grand Bassa, where the ocean is red because of pollution from surface runoff. Southeast of Cape Mesurado are the diabase cliffs at Bassa Point (50 m). The sandy beach continues, with lobes in the lee of nearshore rock outcrops, to Cestos Point, another diabase headland (45 m) near the mouth of the Cestos River. The Sehnkwehn River flows to the coast near Bafu Bay, which is flanked by another headland of Palaeozoic diabase. The surf is disrupted by the many nearshore rocks of diabase along the sandy coast southeast

past Butu River to the Sinoe River. Here, a breakwater has been built to shelter the port of Greenville. Southeast of Greenville, the sandy beaches are interrupted by occasional rocky headlands. Elongated tidal la­­ goons have been formed behind barrier beaches built by the ocean swell, and some river outlets have been deflected many kilometres along the coast (>Fig. 13.6.2). The headlands are bordered by vegetated bluffs rather than active cliffs, but there are many weathering rock outcrops in the nearshore area.

Liberia

Alternations of diabase headlands, sandy beaches, and rocky nearshore areas continue southeast to Fishtown Point and Cape Palmas (30 m), where Russworm Island, on a diabase dyke, was much quarried to provide stone for a causeway to link it to the mainland and provide shelter for the small harbour at Harper, which is used to load timber for export. The coastline then swings eastward across the Ivory Coast border at Cavalla River.

13.6

References Schulze W (1973) A new geography of Liberia. Longman, London Tricart J (1962) Observations on shore morphology at Mamba Point (in French). Erdkunde 16:49–57 White RW (1972) Stratigraphy and structure of basins on the coast of Liberia. Geological survey special paper, 3. Republic of Libe­­ ria,  Ministry of Lands and Mines, Liberian Geological Survey, Monrovia

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13.7  Ivory Coast    Anja Scheffers    

1.  Introduction

descend to low sectors. East of Fresco the coast backed by a Tertiary sedimentary basin and is generally low-lying, with several large elongated lagoons behind a sandy coastal barrier (Pomel 1980). The Ivory Coast has a humid tropical climate with two rainy seasons (May–July and October–November). Abidjan has mean monthly temperatures ranging from 27° in January to 24° in July, and a mean annual rainfall of 1,912 mm. Tides are small, mean spring tide range at Abidjan being 1.0 m. The narrowness of the continental shelf results in a steep submarine slope, which rises too rapidly to produce any substantial reduction of the SSW swell. This great swell originates far away in the South Atlantic where it is generated by the westerly wind belt, and breaks heavily on the coast. It is reinforced in the rainy season by local southwest monsoon winds and is slightly subdued during the January dry season, when it ­encounters

The Ivory Coast extends for about 500 km from the mouth of the Cavally River on the Liberian border to the exit of Aby Lagoon close to Nigeria (>Fig. 13.7.1). The broadly concave south-facing coastline is bordered by a narrow continental shelf: the 120 m isobath, below which the slope plunges to abyssal plains, is only 10 nautical miles from the coast off Abidjan and reaches a maximum of 20 nautical miles off Sassandra. The narrow shelf is interrupted only at one point, off the Vridi canal, where the submarine canyon of the Trou sans fond, which probably follows a fault line, begins immediately offshore. The western part of the coast is generally bordered by bluffs and cliffs at the edge of a coastal plateau on PreCambrian rocks, incised by a series of river valleys that

⊡⊡ Fig. 13.7.1 The Ivory Coast. 1 – Pre-Cambrian shield, 2 – Tertiary sedimentary formations, 3 – Quaternary sands, 4 – Bluffs with a vegetation cover, 5 – Cliffs in Tertiary sandstones west of Fresco, 6 – Low sandy coast, 7 – 120 m isobath, 8 – Direction of swell and angle to the coastline.

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This is a revised version of a chapter on the Ivory Coast by F. Hinschberger (1985) in Bird, E.C.F. and Schwartz, M.L. (eds.) The World’s Coastline. Van Nostrand Reinhold, New York: 585–589. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

13.7

940

Ivory Coast

quickly revegetated. Tricart (1957) has termed these ­intermittently receding tropical cliffs falaises-versants (hillside-cliffs). The coastline at Sassandra is fringed by a narrow sandy beach, which is continuously migrating eastward (Hinschberger and Pomel 1972). The coastline has been modified by human intervention near San Pedro, where a new port was built in 1971 to develop the economy of the southwest Ivory Coast. This area had been virtually uninhabited, but thousands of peasants, forced to leave their lands by the rise of the waters at the Kossou dam (on the Bandama), were moved there. At San Pedro, the river was diverted eastward to prevent the port silting up when sediment was brought down by floods on the San Pedro River. The port was also protected from accretion of wave-deposited sand drifting alongshore by a breakwater on the western side and artificial groynes. Eastward drifting sand was trapped in a prograded sector, permitting the reclamation of several square kilometres for housing development. Near Fresco, the Pre-Cambrian basement disappears under Tertiary strata, consisting mainly of soft sandstones, which have been cut back as vertical cliffs that extend for a distance of several kilometres. Between Fresco and Abidjan, the angle of the swell in relation to the coastline is greater than in the west (60°), but the resultant longshore drifting is still fairly strong, supplying sand to a long, straight, and mostly narrow Holocene coastal barrier. This barrier stands in front of a former embayed coast (>Fig. 13.7.2), enclosing a series of elongated coastal lagoons, the outlines of which may be partly due to tectonic subsidence (Tastet 1977). It may have been initiated on this alignment by emergence, perhaps during a midHolocene phase when sea level fell relative to the land, but longshore spit growth may also have contributed. The barrier carries rich scrub and forest vegetation. In

the northeast wind, locally called the harmattan. The swell gives rise to incessant large waves known in French as the barre and in African dialect kalema. These produce steeply sloping beaches of coarse sand, on which giant beach cusps are usually present. Because of the curvature of the coastline, the angle of incidence of the SSW swell varies alongshore. From Cape Palmas to Fresco, it arrives at the coastline at an angle of 45°, which results in strong eastward longshore drifting but further east the angle increases to 60° and more, and the resultant longshore drifting is less strong.

2.  The Ivory Coast East from the Liberian border, the Pre-Cambrian shield comes to the coast. There are moderately high cliffs and bluffs, the elevation of the hinterland only occasionally exceeding 100 m. The shores are rocky, with sandy beaches drifting eastward. The coastline is set back in steps related to a series of echelon faults that divide it into seven sectors. One of these led to the formation of the Grand Berebi cove; another to the beautiful Monogaga cove, between San Pedro and Sassandra. Although the main irregularities of the coast were tectonic in origin, the outlines of the dozens of small bays and miniscule capes that separate them is related to the lithology of outcropping rocks. The capes correspond to the outcrops of hard rocks, often veins, the bays to less resistant rocks, such as weathered granites. The steep coast consists mainly of bluffs with a dense vegetation cover, and  only on the capes is the country rock exposed as a basal cliff, cut by strong wave action. The bluffs are not abandoned cliffs, however: they are receding as the result of intermittent mass movement, forming scars, which are

⊡⊡ Fig. 13.7.2 The coast near Abidjan, showing the Vridi Canal, cut through the barrier from the Ebrié Lagoon and the coastline changes since it was cut in 1950. Lagoon

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the Grand Lahou Lagoon and the Ebrié Lagoon west of Abidjan, there are elongated sandy islands that may be the dissected remnants of an older, possibly Pleistocene, inner barrier. The Bandama River flows through a large swamp into the eastern end of the Grand Lahou Lagoon, and out through a natural gap in the coastal barrier. The beach is interrupted, but a looped sand bar has formed off this gap, produced by the interaction of outflowing currents from the river and lagoon, which tend to wash the sand away, with incoming ocean swell, which breaks across the bar and tends to wash the sand back into the gap in the coastal barrier. Sand drifting alongshore by-passes the Grand Lahou outlet along this looped sand bar. The making of a port at Abidjan required access to and from the sea through the coastal barrier. The first attempt to achieve this was in 1905, when a canal was cut through a narrow section of the barrier at Port Bouet, but this was immediately blocked by sand drifting eastward and proved useless. In 1950, a second attempt was more successful. A canal was cut through the coastal barrier at Vridi, and a breakwater built on the western side to prevent sand drifting across it. Sand drifting from the west accumulated alongside this breakwater, and formed a long submarine bar off the entrance, which ships must circumnavigate to gain access to the canal. This bar disappears near the Trou sans fond, and sand flows down into the submarine canyon. Interception of drifting sand has prograded the coastline to the west of the breakwater but left the coast unprotected immediately to the east. As a result, over a kilometre to the east, an embayment has been cut back into the coastal barrier, breaching the coast road (2). The opening of this permanent artificial canal to the sea has led to an increase in the salinity of the Ebrié Lagoon

13.7

and modified its flora and fauna. The Comoe River now flows into the eastern end of the Ebrié Lagoon, its natural mouth at Grand Bassani, which was permanently open until 1950 but is now blocked by a sand barrier, which is cleared for only two or three months of the year, at times of floodwater discharge. From Abidjan to the Ghanaian border, the coastal barrier swings ESE, so that waves arrive more or less at right angles to the coastline, and generate very little longshore drifting to the east. On occasion, the less frequent southerly swells generate westward longshore drifting. This is a sector where drifting sand has accumulated, so that the coastal barrier is wide, attaining several kilometres east of Grand Bassani, with numerous parallel beach and dune ridges. Towards Assinie, the barrier narrows again in front of the large Aby Lagoon, which extends into a valley where tectonic subsidence has occurred (Tastet 1977). In the absence of longshore drifting, the beaches east of Abidjan show circulating cells of water and sand. Ocean swell moves sand on to the beach but the breaking waves generate rip currents that return it to the sea floor. These cells are clearly visible on aerial photographs, particularly off Assinie beach, where they interrupt the lines of foam.

References Hinschberger F, Pomel R (1972) The morphology of rocky coasts between Monogaga and Sassandra (Ivory Coast) (in French). Ann Univ Abidjan G4:7–37 Pomel R (1980) Physical geography of the lower Ivory Coast (in French). Norois 27:348–350 Tastet JP (1977) The quaternary of the Ivory Coast (in French). Recher­ ches  Françaises sur le Quaternaire. Bull Assoc Fr Etude Quat, Suppl:172–186 Tricart J (1957) Geomorphological problems of the western Ivory Coast (in French). Bull Inst Fr Afr Noire 19:1–20

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13.8 Ghana

L.A. Dei

1.  Introduction The Ghanaian coastline is about 600 km long, of which 280 km are sandy beaches. The rest is steep, partly cliffed, with rocky hores interrupted by sandy beaches in bays. The area is underlain by Pre-Cambrian and Palaeozoic rocks, and isolated outcrops of Devonian sandstones occur around Accra and between Cape Coast and Takoradi. The coast of Ghana has an equatorial climate with two rainy seasons (May–July and October–November). Accra has mean monthly temperatures ranging from 27° in January to 25° in July, and a mean annual rainfall of 1,402 mm. Tides are small, mean spring tide range at Takoradi and Accra being 1.3 m. The coast receives a strong SSW swell that originates down in the South Atlantic where it is generated by westerly gales. It breaks heavily on the coast, reinforced in the rainy seasons by local southwest monsoon winds and slightly subdued during January–April, when it is opposed by the northeast wind, locally called the harmattan. The swell breaks into surf that produces steeply sloping beaches of generally coarse sand on which giant beach cusps are often present. The beaches border cliffs and bluffs or lagoons and swamps, but dunes are poorly developed, largely because the beach sand is generally moist, inhibiting wind transport. An exception is at Old Ningo where the sand is unusually coarse and shelly, and has dried out sufficiently to be blown to backshore dunes. Evidence of higher sea levels in Pleistocene times is in the form of a stairway of three beach terraces (Dei 1972a). Terrace I (between 9 and 12 m above present sea level) is presumably the oldest beach terrace. Sieved sand samples show a characteristically asymmetrical sigmoid curve, made up of 4.0% coarse sand, 15.5% medium sand, 49.8% fine sand, 12.8% silt, 15.2% clay, and 2.7% organic matter. It is moderately sorted, with sorting values ranging between 1.7 and 2.8. Particle surface characteristics show that most grains are matt, with ferruginization largely in cavities. Quartz particles are mostly irregular, between 200 and 250 µm or below, but subrounded between 400

and 630 µm or above. These beach sediment were derived mainly from continental deposits that had been partially reworked by wave action. Terrace II (between 4.5 and 9 m above present sea level) is more extensive than Terrace I, and separated from it by a well-defined slope varying between 10 and 15°. Sieved samples of quartz grains show perfect sigmoid curves, with average sorting values between 1.29 and 1.34. Quartz particles have higher roundness indices than those of Terrace I. Surface texture remains matt, as in Terrace I, but there is a significant increase in the proportion of grains with polished surfaces, perhaps because of successive reworking of older beach materials. Terrace III (less than 4.5 m above present sea level) is the most recent terrace along the Ghanaian coastline. Samples show perfect sigmoid curves, and quartz particles are more rounded but less ferruginized than those of Terraces I and II. Seaward terminations of coastal streams are occupied by lagoons of various shapes, some finger-like, others pear-shaped. A striking feature of the lagoons is the concentration of calcareous nodules about 0.3 m below the surface of former lagoon shores. Their shapes are irregular, resembling the roots and stems of plants that are extinct in the area: they are phytoconcretions. Fossil shells are found encrusted in these nodules, the shelly material being the source of the cementing calcite. Cliffs and steep vegetated bluffs descend to rocky shores, with some beachfringed sectors (Dei 1975).

2.  The Coastline of Ghana From the Ivory Coast border, the coast first trends slightly south of east, and is lined by sandy beaches and barriers. The mouth of the Ankobra River has been deflected eastward by the predominant longshore drifting. Near Axim, Pepre Point has emerged beach terraces at 5 and 12 m above low tide level, which contain rounded pebbles of the Upper Birrimian. Mounds with angular or semi-rounded boulders mark the positions of former islands. Pieces of

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rounded and angular beach rock occur, and coffee rock (iron-cemented beach sand) is seen in various stages of consolidation. At Prince’s Town, a granite promontory rises 30 m above sea level, with an emerged shore platform about 12 m above present sea level. Active cliffs rise to about 15 m above sea level. At Dixcove, there are no shore platforms, but islands and scattered angular granite or granodiorite boulders. Toward the east of the town, a pebbly beach terrace consisting of semi-weathered granite pebbles is seen about 9 m above low tide level. This terrace, which overlies the boulder zone and is ferruginised, indicates a higher (probably Pleistocene) sea stand. At Takoradi, there are beach terraces up to 15 m above sea level. Erosion is in progress along cliffs and beaches. Beach rock occurs alongside coffee rock on eroding beaches and is generally fragmented, but outcrops of it overlie the feldspathic Elmina Sandstone. Incipient beach rock is seen in many places. Much of the coast of Sekondi has been fortified against erosion by the construction of sea defences. Beach erosion is seen at Turtle Cove east of Sekondi. At least three abrasion platforms can be identified in sedimentary formations of varying hardness, and there is evidence of a collapsed cave roof. Cliff notches and stacks also occur. At Gold Hill (Komenda), a natural cave about a metre above high tide level penetrates the hard limonitic face of a cliff, and opens into a softer leached zone. There is a blowhole, overgrown with vegetation, about 10 m from the cave entrance. To the west is an emerged beach 1 m above high tide level, which has been colonized by herbaceous plants. Behind the beach and to the east of the cave are bluffs almost completely covered by vegetation: they were former cliffs, degraded by subaerial processes after emergence. At Elmina, the feldspathic Elmina Sandstone outcrops, with surfaces pitted by circular holes and polygonal patterns. Marine erosion has penetrated along joints, forming rectilinear inlets, and abrasion platforms are found 0–1 m, 1–2 m, and 3 m above present sea level. Much of the coast between Cape Coast and Abandzi, near Saltpond, has abrasion benches 12 m and 1–4 m above low tide level. The estuary of the Ochi (Amisa), 5 km east of Saltpond, is littered with rounded pebbles, reworked material from the Amisian formation. At Apam rounded outcrops of metamorphosed lava are seen, and small cliff base notches are common. At Winneba, the beach is rich in calcareous sediment. There is slight induration of beach sand but no beach rock

has yet formed. Cliffs are generally low, except to the west of the town, where there are abrasion platforms 10–12, 3, 1–2, and 0–1 m above present sea level. Senya Beraku has three main cliffs: a high (30 m) cliff, an intermediate cliff (10–15 m high), and a low (6 m) cliff. There are also abrasion benches 15, 10–12, 3, 1–2, and 0–1 m above present sea level. Caves and notches have been cut into the cliffs, and beach rock outcrops at their base. Iron-cemented beach sand and gravel has been preserved on a rock bench at 13–17 m above sea level at Accra (McCallien 1962). The same terrace with a cliff between 14 and 15 m is seen at Tema (Davies 1964). Emerged beaches between 1.1 and 8 m above high tide level occur between Accra and Tema (Dei 1975). At the mouth of the Sakumo Lagoon, near Tema, is a beach rock outcrop about a metre above mean sea level (Dei 1972b). Beach rock also outcrops at Prampram, where at low tide, it can be traced seaward for a distance of at least 40 m. East of Accra is the lobate delta of the Volta River, its seaward fringe a sandy coastal barrier interrupted at the river mouth and backed by lagoons (Songaw, Angaw, and Keta Lagoons) and swamps, which occupy much of the delta plain. The Quaternary delta sediment are marine, fluviomarine, and fluvial, and occur down to a depth of 170 m below sea level, underlain by Cretaceous located at over 200 m (Ghana Public Works Department 1960), showing that there has thus been Quaternary subsidence in the Volta delta region. The absence of surface features indicating Holocene progradation, such as beach ridges, implies that this subsidence has continued. To the north is the early Holocene coastline, rising to a land area with red lateritic material formed by prolonged tropical weathering of the Tertiary rock formations. Offshore contours indicate that the sea floor declines steeply, dissected by a number of submarine valleys, including a major canyon off the Volta River mouth. The steep seaward margin of the delta is a monoclinal flexure, a feature quite common along the Ghanaian shoreline. The delta coastline has long been subject to intense marine erosion, which has attracted both national and international attention. Records indicate that beach erosion here dates back at least to the beginning of the twentieth century. There is an increase in wave energy from west to east (Dei 1975), but the erosion is probably the outcome of subsidence in the delta region. Some progradation has occurred on spits bordering the river mouth, but this is largely local accretion of sand eroded from other parts of the delta coastline and carried to the river mouth by longshore drifting.

Ghana

References Davies O (1964) The Quaternary in the coastlands of Guinea. Jackson Son & Co., Glasgow Dei LA (1972a) The central coastal plains of Ghana: a morphological and sedimentological study. Z Geomorphol 16:415–431 Dei LA (1972b) Some observations on Ghanaian beach rocks. J Trop Geog 35:26–31

13.8

Dei LA (1975) Morphology of the rocky shoreline of Ghana: Eustatic? Bull Ghana Geogr Assoc 17:1–30 Ghana Public Works Department (1960) Report on coastal erosion, Sea defences and Lagoon flooding at Keta. Ghana Public Works Department, Accra McCallien WJ (1962) The rocks of Accra. Curwen Press, London

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13.9  Togo and Benin

André Guilcher

Introduction The coastlines of the two small West African states, Togo and Benin, are similar and can be discussed together. Togo has an 80 km coastline and Benin about 110 km. The large West African Pre-Cambrian shield slopes beneath a coastal belt of Cretaceous, Tertiary, and Quaternary rocks 35–130 km wide. The Upper Tertiary sedimentary capping rocks consist of continental sandy clays, called the Terre de Barre (barral means clay in Portuguese), which form a plateau 20–70 m above sea level. This plateau was deeply incised by rivers during Pleistocene low sea level phases, but the valleys thus formed have been partly filled by Late Pleistocene and Holocene sediment that form the coastal complex, bounded along the Atlantic shore by a long, straight, monotonous sandy barrier beach which impounds a chain of lagoons and swamps. The Togo and Benin has a humid tropical climate with two rainy seasons (May–July and October–November) and a relatively dry season from December to May. Lomé has mean monthly temperatures ranging from 27° in January to 24° in July, and a mean annual rainfall of 875 mm. Tides are small, mean spring tide range at Lomé being 1.4 m. The coastal dynamics are generated by the surf which breaks continuously on the shore as a consequence of swell arriving from the SSW, originating in the storm belt of the Southern Ocean. The continental shelf is narrow (about 20 km wide) and the ocean wave energy remains high when the surf reaches the shore. The swell varies in height with conditions in the generating area, and is often stronger in mid-year, the southern winter. It generates eastward drifting of the order of 1 to 1.5 million cubic metres/year. Local winds are generally light, and blow from the southwest. They have little effect on the shore surf, but have influenced the evolution of inner lagoonal shores. Another factor in coastal evolution is river discharge, which differs with the seasons, decreasing steadily from November to May. The Mono and Oueme Rivers show one

prominent peak in September while the Couffo River has two peaks in June and October, with a secondary low in August (Guilcher 1959). The sand forming the beach drifts along the shore to the east because the swell breaks obliquely on the coasts of Togo and Benin, as it does on the entire coast of the Gulf of Guinea. It is a quartzose sand, which is supplied from two sources to the west: rivers, the larger one being the Volta in Ghana, and eroding cliffs, which exist in a number of places in Ghana and on the Ivory Coast, and from the sea floor. In recent decades the longshore sand supply has decreased, and the beach-fringed coastline has retreated. This is partly because of the building of large dams across rivers, which have intercepted the fluvial sand, and the building of breakwaters to protect harbours, which have trapped sand on their western sides. Every state wants at least one dam and one harbour, sometimes more, and that has been the problem. In 1967 a deep-water port constructed in Togo also disrupted longshore drifting and formed a new sediment cell (Anthony and Blivi 1999). In Benin the extraction of beach sand for construction purposes, promoted by the government as resource use and providing employment, has resulted in backshore erosion and the destruction of buildings, including the homes of some of those employed. The outer beach borders a coastal sand barrier built in Holocene times. Where it is wide enough, the barrier includes several parallel beach ridges, as in the Anecho region in Togo), and between the barrier and the plateau are extensive lagoons and swamps (> Fig. 13.9.1). The mouths of rivers and the outlets from lagoons are open when the outflow is strong, particularly in wet seasons, and closed when it slackens, and waves are able to build up sand across them. The location of these outlets tends to migrate eastward along the coast in response to longshore spit growth, and then to return suddenly westward as the result of the breaching of a new outlet. In western Benin, for example the outlet was situated at Boca del Rio in 1954–1955, and closed when waves built up sand across it during the 1955–1956 dry season. At the beginning of the wet season in October the outlet remained

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⊡⊡Fig. 13.9.1 Swell breaking on the beach in Western Benin, showing the coastal barriers, lagoons, swamps and an inflowing river.

closed and water level rose to submerge the backing swamps. Then the local authorities ordered the cutting of  a channel across the barrier, forming an outlet that remained open, even during the following dry season. Outflow then acted as a natural breakwater, so that the beach widened on the western side. The lagoons behind the coastal barriers are only about a metre deep, except where they include the channels of rivers crossing to outlets (Guilcher 1959). The lagoon floors are muddy, but sandy beaches exist on their shores, on which the waves generated by the local southwesterly winds build small hooked spits, evenly spaced, tied to the mainland in the west and recurved at their eastern ends. Behind the coastal lagoons there are extensive swamps that contain many sandy strips elongated from west to east and much smaller sandy islands or mounds that are more or less circular in plan. The elongated sand deposits are remnants of earlier (Pleistocene) coastal barriers that were dissected by incised rivers and creeks during low sea level stages, then partly submerged and surrounded by swamps. The subcircular sandy mounds, which are prominent in the area between Grand Popo and Ouidah, are remnants of former beach ridges that were dissected by the meandering of creeks and the shifting courses of rivers. The various deposits in the coastal region have resulted in the straightening of the coastline, which was previously

much more indented. Lake Ahéméand other lakes in the lower reaches of the rivers are former rias, drowned river valleys that became separated from the sea by the for­ mation of coastal barriers during and after the Holo­ cene  marine transgression. In Lake Nokoué salinity has increased and there have been ecological changes fol­ lowing the opening of a marine entrance at the port of Cotonou. The Benin marshes have been exploited for salt production, which has resulted in numerous small ponds in the Grand Popo area. This is an example of anthropogeomorphology. Another result of this salt extraction has been the extensive cutting of the mangrove trees for firewood. Mangroves have become sparse, and have been replaced by marshes and mudflats, even though there is no long dry season.

References Anthony EJ, Blivi AB (1999) Morphosedimentary evolution of a deltasourced drift-aligned sand barrier-lagoon complex, western Bight of Benin. Mar Geol 158:161–176 Guilcher A (1959) Coastal sand ridges and marshes and their continental environment near Grand Popo and Ouidah, Dahomey. Proceedings of the 2nd coastal geography conference, Office of Naval Research, Washington, D.C., pp 155–167

13.10  Nigeria

Etop Usoro

1.  Introduction The coast of Nigeria is about 850 km long. It has a hot, wet climate, with an annual rainfall of at least 3,500 mm. There are two rainy seasons (May–July and October– November) and a relatively dry season (NE trade winds) from December to May. During the wet seasons moistureladen southwest winds from the Atlantic Ocean produce copious rainfall, which makes places like Calabar, Bonny Town and Port Harcourt among the rainiest in the world. Temperature ranges from 20°C to 35°C, and humidity is persistently high, ranging from 70% at Lagos to over 80% at Calabar. Mean spring tide range increases eastward from Lagos (1.0 m) to the estuaries of the Imo and Cross rivers (2.6 m and 3.0 m respectively). Brackish water spreads far inland, extending to the apex of the Niger delta near Onitsha, about 220 km from the open ocean. Bonny Town on the east coast of the Niger delta has a tide range of 1.9 m. The Nigerian coast is characterised by strong wave action. The strong southwest monsoon is the dominant wind, and blows all year round. The open Atlantic provides a long fetch, producing powerful waves, generated by these prevailing winds. Longshore currents set up by the breaking of these waves distribute sediment along the coast. In the western sector of the Nigerian coast there is a strong west-to-east longshore drift. At the Lagos Harbour mouth, the longshore drift has been interrupted for many years by a breakwater, which has trapped sand moving eastward along the coast. The geology is dominated by the crystalline base­ ment complex and folded Palaeozoic strata of West Africa, which pass beneath thick recent sedimentary accumu­ lations towards the coast. The Nigerian coastal plain is defined by a massive sedimentary basin about 350 km wide and 670 km long (Allen 1965a). The sedimentary formations dip gently seaward. The Holocene sediment, consisting of fluvial and lagoonal deposits and littoral sands, extend inland to the apex of the delta near Onitsha and the lower valleys of the Ogun, Shasha, Imo, and Cross rivers. A marine transgression that deposited older quartzose sand was followed

by ­depositional phase from the late Pleistocene to the early Holocene, during which the silts, clays and sands of the modern delta and the associated barrier islands and mangrove swamp environments were deposited (Allen 1965b). The coast consists of a low-lying depositional plain characterised by extensive estuarine lagoons and mangrove swamps fronted by beach-ridge barriers. The Niger is the largest of several rivers that transport enormous quantities of sediment of all grades to the Gulf of Guinea, where waves, tides, and currents sort and redistribute the materials and build them into a series of sedimentary formations. The basins of the Niger and Benue yield about 60% of the coastal sands, the remainder coming from the basins of the other Nigerian rivers. The rivers carry mostly medium to coarse sand, their characteristic heavy minerals being green and brown hornblende, epidote, zircon, and staurolite. Apart from sand, these rivers transport vast amounts of silt and clay in suspension during flood discharge to the coast. The deposition of such organic-rich clayey silt and plant debris produces extensive mudflats and swamps. The floors of the large creeks and tidal channels in the delta contain thick sand beds, while many of the smaller channels contain sand interbedded with dark clayey silt and plant debris. During the Last Glacial low stand of sea level the rivers became incised and extended across the conti­nental shelf (Allen 1965b). A transgressing sea advanced inland to deposit quartzose littoral sands. Once the sea level became relatively stable, progradation occurred forming beachridge barriers, mangrove swamps, organic-rich estuarine tidal flats, and river flood plains and deltas. The advance of the Nigerian coastline has now given place to coastal retreat, with widespread truncation of the seaward margins of barrier beaches. Coastal vegetation includes forest extending on to steep  coastal slopes, scrub and woodland with stunted palms on dunes and beach ridges, and mangroves in estuaries and lagoons. The mangroves of the Nigerian coast are of two types. Swamps fringing lagoons and estuaries are dominated by the stilt-rooted red mangrove (Rhizophora

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r­ acemosa), backed by Avicennia with pneumatophores. Freshwater swamp vegetation gradually displaces the mangroves farther inland, and on more elevated land flood plains and coastal terraces rain forest becomes dominant. Mangrove swamps are extensive along the Nigerian coast. Within the Niger delta and its flank areas, the swales between sand barriers are occupied by mangrove swamps, and the estuaries of the Imo, Kwa and other rivers are bordered by mangroves. They also occur on lagoon shores, as on the northern shores of Lekki Lagoon. The mangrove swamps occupy tidal mudflats laced with tidal channels, a myriad of meandering creeks that vary in size and shape, from widths of more than 1 km to small winding waterways about 20 m wide. At their seaward ends, the channels broaden into large tidal passes more than 2 km wide, and some are studded with complexes of shoals and mudflats. Intricately connected creeks can be seen in the vicinity of Port Harcourt, where a complex network of drainage channels has broken the mangrove swamp into irregularly shaped areas. The mangrove swamp environment along the Nigerian coast forms at an unusually rapid rate. On many parts of the Niger delta mangroves are advancing rapidly. Mean­ dering of tidal creeks enables the mangroves to invade former sandy barriers, and remnants of the sandy formations form higher ground where settlements and farms are located. Mangrove swamps are succeeded landward by firm alluvial ground, or a depositional terrace constructed by the accretion of sand during flooding. This is the coastal flood plain, which occupies a broad space between the mangrove swamp and the sandstone terrace. Large meandering scrolls enclosing ridged point bars, low natural levees that descend into heavily forested backswamps, and cutoff channels and oxbow lakes are prominent features of the landscape. The levees vary in height between 4.5 and 8 m above the backswamps, and where they are breached by floodwaters, large crevasses have formed. The backswamps are seasonally flooded, and settlements, farms and lines of communication are restricted to the levees.

2.  The Nigerian Coastline East from the Benin border, between Badagry and Mahin, the south-facing coast consists of sandy beaches and barriers backed by lagoons. The outer barrier beaches consist largely of clean, coarse, angular sand, although medium and very coarse grains are seen as well. The inner barriers have predominantly finer and more rounded grains than do the outer barrier beaches. The high degree of ­roundness

of the inner sand ridges may be attributed either to the changes that have affected them between their derivation and deposition or to post-depositional physicochemical changes. Silt and some organic debris have been deposited in the swales between the parallel sand ridges. Along Victoria Beach, Lagos, coarse quartzose sand is found on the upper parts of the beach while finer sand characterises the foreshore, which also contains a high proportion of calcareous material. West of the Niger delta, mud pebbles are found on eroding stretches of the beach. Sandy barrier beaches constitute the most striking ­features along the Nigerian coast. They are backed by a low degraded cliff about 10 km from the shore that rises to the margin of the coastal terrace or plateau. Between the terrace and the sea, a depositional plain includes a series of  long sand ridges trending parallel to the coastline (>Fig.  13.10.1). The sand ridges are separated from one another by elongated muddy depressions, giving a ridge and swale topography. The sand ridges may be classified into a younger outer group, formed during the Holocene, and including active beach ridges, and an older inner group of Pleistocene age. There are two sets of beach ridges on the outer barrier near Badagry: one set runs diagonally across the landward side of the barrier, and the other set is aligned parallel with the coastline. The beach ridges are no more than 2 m high and vary in width between 40 and 50 m crest to crest. Near Mahin some beach ridges can be traced for distances of several kilometres without interruption. The outer barrier is separated from the inner sand ridges by creeks and meandering channels. West of Lagos the lagoon subdivides into several smaller creeks to form an intricate pattern of waterways, but east of the city it broadens into an expansive water body attaining a maximum width of about 7 km in the Lekki Lagoon. Lagos Harbour is located in the lagoon near the only connection with the sea. The barrier-lagoon complex of western Nigeria constitutes a system in which the parallel sand ridges mark lines of growth, each having once been a beach, separated from the previous one by a lagoon. Some of the lagoons eventually silted up as a result of sedimentation to form swampy depressions. Badagry Creek, which continues east of Lagos as the Lekki Lagoon, is the last such enclosed portion of the sea not yet filled with estuarine sediment. The sub­ sequent development of barriers along the full length of the coast in this area led to a marked straightening and shortening of a formerly indented and embayed coast, as indicated by the margin of the coastal terrace (Guilcher 1959). West of the Niger delta, the coast has a remarkable display of a barrier beach–lagoon complex in which parallel

Nigeria

sand ridges are spaced between filled depressions and open creeks. The Lagos and Lekki Lagoon, which has provided Nigeria with a magnificent sheltered harbour, is separated from the open sea by a broad sand barrier with a wide surf-bound beach. The central portion of the Nigerian coast is dominated by the Niger delta, which consists of a thick pile of sediment, much of which has accumulated since Cretaceous times in a down-warped coastal margin (Allen 1965a). East of the Niger delta, the coastal landscape is typically that of beach ridges spaced between the estuaries of the Imo and Kwa rivers, while much of the Cross River estuary is flanked by mudflats and mangrove swamps. The Nigerian coastal plain indicates a long period of seaward growth and represents an excellent example of a depositional regression of the sea by littoral sedimentation. At the apex of the Niger delta there is an eastward current, which continues to the Cross River estuary, and a westward current on the flank of the delta to the mouth of the Benin River. These currents drive sand brought by the rivers northwestward in the western delta and eastward between Bonny and the Cross River estuary. This littoral drifting is the most important agent moulding sandy shoreline features along the Nigerian coast. Tidal flows are also important for sediment dispersal, especially in the mangrove belt and nearshore waters.

13.10

Access to swamps and mudflats is through the mouths of rivers and major distributaries of the Niger. In the Lagos area, the tide penetrates the lagoon through the harbour entrance. Tidal flows are diurnal and create reversing currents. Tidal currents in the swamp belts of the Niger delta and the lagoon complex west of the Benin River are not consistent because of the intricate geometry and numerous interconnections of the creeks within the mangrove swamps. The maximum tidal velocities in the mangrove swamps vary from 40 to 180 cm/s. Such velocities are strong enough to shift coarse sand. The dispersive powers of the tide diminish with increasing depth and distance from the shore. East of the Imo River estuary, the tidal flats and mangrove swamps that characterise much of the Niger delta disappear from the coastal landscape, giving way to a complex strand-plain barrier up to 10 km wide. This barrier consists of numerous sub-parallel sand ridges formed by successive additions of beach ridges to the coastline (>Fig.  13.10.1). The intervening swamp hollows are reflected in the strong lineation of vegetation of dense raffia thicket, which differs markedly from the sparse scrub on the sand ridges. The summit levels of the ridges decrease seaward from 7 m near the margin of the sandstone terrace to 3 m near the coast. Some of the successive sand ridges may have developed from spits growing eastwards

⊡⊡ Fig. 13.10.1 Beach ridges on the Nigerian coast east of the Imo River estuary. (Courtesy Geostudies.)

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with the longshore sand drift. One such spit can be seen west of the estuary of the Kwa River. To the north of this spit, the vegetation pattern reveals the incorporation of former hooked spits. Near the sandstone terrace, the ridges in the strand plain are of greater elevation and may have been longshore bars that emerged above sea level. Near the Cross River estuary, sets of older beach ridges are truncated by newer ones. The lines of dis­ cordance between groups of sand ridges indicate that periods of coastal progradation alternated with periods of erosion, presumably relating to variations in sediment supply (Allen 1965a). The Nigerian coast is currently receding, possibly following a slight rise in sea level. Everywhere along the coast, sand cliffs characterise the backshore, and arid beaches are being severely eroded. Between Lagos and the Benin border near Kweme the sandy barrier has been cut back rapidly, at a rate of more than 5 m a year. Severe erosion is also occurring along the shore of barrier islands fringing the Niger delta, especially between Bonny and Andoni. Near Lagos the trapping of longshore drift on the western side of the breakwater built to maintain the port entrance at Lighthouse Beach has caused accretion,

accompanied by rapid recession of Victoria Beach to the east, where attempts have been made to restore it by beach nourishment with sand pumped in from the sea floor. In recent years erosion has been severe during recurrent storm surges (Olaniyan and Afiesimama 2003). Discordance between sets of beach ridges is indicative of coastline changes. At Mahin, for example, older beach ridges have a trend that differs by more than 25°C from the younger ones. East of the Kwa River estuary three distinct episodes of coastal accretion with intervening periods of retreat have been identified.

References Allen JRL (1965a) Late Quaternary Niger delta and adjacent areas. Bull Am Assoc Petrol Geol 49:547–600 Allen JRL (1965b) Coastal geomorphology of eastern Nigeria beach ridge barriers and vegetated tidal flats. Geol Mijnbouw 27:1–21 Guilcher A (1959) Coastal sand ridges and marshes and their environment near Grand Popo and Quidah Dahomey. Proceedings of the 2nd coastal geography conference, Washington, pp 189–212 Olaniyan E, Afiesimama EA (2003) Understanding ocean surges: A case study of the Victoria Island bar beach Lagos. African Marine Science

13.11 Cameroon and Equatorial Guinea

Maurice Schwartz

1.  Introduction The coasts of Cameroon (>Fig.  13.11.1) and Equatorial Guinea (>Fig. 13.11.2) share a very similar physical environment. They lie along a trailing-edge coast where PreCambrian basement rocks (gneiss, schist, quartzite, and migmatite) outcrop, mantled in places by Cainozoic deposits (Orme 2005). The eastern edge of the Niger Basin extends into Cameroon. Recent volcanic activity has contributed to the evolution of the coastal landscape. The coast of Cameroon lies between 2°N and 14°N and that of Equatorial Guinea between 1°N and 2°N. The climate is equatorial with a SW monsoon influence. In the region of Douala, mean annual rainfall is just over 4,000 mm, with temperatures between 24°C (July– August) and 27°C (January–March): there are four seasons: rainy from March to June and again from September

to October, a brief dry spell in August, and a long dry season from November to March (Van Chi-Bonnardel 1973). Average runoff along the whole sector is greater than 1 m/year. In these warm and wet conditions, chemical weathering proceeds rapidly and soils in the coastal regions are either latosolic or lateritic (Pritchard 1971). South-westerly monsoon winds prevail, especially in mid-year, and waves arriving from this direction accompany a southerly swell that originated in the southern Atlantic Ocean, and moves in on either side of the high island of Fernando Poo. The highest swell coincides with the strongest and most frequent onshore winds, from June to September. Mean wave heights over 1.5 m occur less than 20% of the year (Orme 2005). Tides are semidiurnal and with a mean spring tide range of about 2 m: Douala has 2.1 m and Kribi 1.6 m.

⊡⊡ Fig. 13.11.1 Location map and geomorphology of the Cameroon coast.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.11, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Cameroon and Equatorial Guinea

⊡⊡ Fig. 13.11.2 Location map and geomorphology of the Equatorial Guinea coast.

2.  The Cameroon Coastline The coastline of Cameroon is about 200 km long, bordering the Gulf of Guinea. East from the Nigerian border, it is low-lying and fringed by mangrove swamps and intermittent sandy beaches, continuing the deltaic coastal plain of Calabar. The hinterland then steepens on the slopes of the active volcano of Mount Cameroon (4,095 m), and between Idenao and Victoria there are rocky shores, with some cliffed sectors running out to headlands such as Cape Nachtigal (Ngwa 1967). East of the volcano, the coast is again low-lying, the lowlands widening to more than 100 km behind the mangrove swamps and low islands around Douala, where the Wouri and Dilamba rivers flow into an estuarine lagoon bordered by a barrier spit on the eastern side. Douala is a major seaport. Sandy beaches extend along the coast south to the Sanaga River delta, which has spits at the river mouth resulting from northward longshore drifting of sand by waves arriving from the south and southwest.

South of the Sanaga delta, the coast curves to face westward, and remains low-lying south past Kribi, a small seaport. The beach-fringed coast is interrupted by many inlets, creeks, and lagoons with sand bars and spits at their mouths and mangrove-fringed shores and islands. South of Kribi, there are some shore and nearshore coral reefs, but these come to an end near the mouth of the River Niem at Cambo.

3.  The Coastline of Equatorial Guinea The mainland coast of Equatorial Guinea (Rio Mbini), about 150 km long, is low-lying, with intermittent sandy beaches and mangroves. The coastal plain rises inland in a series of Pleistocene terraces, and hilly ranges come to the coast southward near Cabo San Juan. Bata is a port on the northern coast, and the Mbini River, joined by the Lana River, flows into an elongated estuarine inlet at Mbini while the River Temboni and Noya flow into a similar inlet southeast of Cabo San Juan.

Cameroon and Equatorial Guinea  

   13.11

Longshore drifting is northward, but is not strong: there are no major spits or river-mouth deflections. Southwest of Mbini, there are coral reefs extending past Punta Yoni. Corisco and Elobey are small low islands off the south coast of Equatorial Guinea.

slopes of this island are intensively farmed, but the southern part is rugged and dissected by rivers that descend cataracts. Along the southern coast of the island, exposed to the prevailing SW swell, is a coral reef, and the sandy beaches include coralline as well as basaltic sand and gravel. 

4.  Fernando Poo (Bioko)

References

Northwest of Equatorial Guinea in the Bight of Biafra is Fernando Poo, now known as Bioko, a mountainous island with an area of 2,034 sq. km (Van Chi-Bonnardel 1973). It is a volcanic island with three basaltic cones, a southwest extension of the volcanic zone on the Cameroon coast. The highest elevation is 3,007 m at Pico Santa Isabel. The steep

Ngwa JA (1967) An outline geography of the Republic of Cameroon. Longmans, London Orme AR (2005) Africa, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 9–21 Pritchard JM (1971) Africa: The geography of a changing continent. Africana Publishing Corp., New York, p 248 Van Chi-Bonnardel R (1973) The atlas of Africa. Jeunse Afrique, Paris

955

13.12 Gabon, Congo, Cabinda and Zaïre

P. Giresse

1.  Introduction The combined coastline of Gabon, Congo, Cabinda (a separate province of Angola) and Zaïre (now the Democratic Republic of Congo) is about 980 km long, extending from 1°N of the Equator to 6°S (Giresse and Kouyoumontzakis 1985). It forms the oceanic boundary of two Mesozoic and Tertiary sedimentary basins, those of Gabon and the Congo, which are separated by the Mayumba salient of the Pre-Cambrian basement (>Figs. 13.12.1 and >13.12.2). The land surface is mantled by the Serie des Cirques, a piedmont deposit of silty clays, gravels and laterites, probably of Plio-Pleistocene age, but the Upper Cretaceous rocks outcrop along the coast of southern Congo. Differential erosion of outcrops on either side of fault lines is responsible for the outline of the successive Pointes: Pointe Pedras, Pointe Banda, Pointe Kounda, Pointe Indienne, Pointe Noire, and Pointe de Tafe. At Pointe Noire one of these faults brings resistant dolomites on the headland against uncemented sands and clayey marls, cut out to form the adjacent bay (Giresse and Kouyoumontzakis 1971). This coast is generally considered to have been stable, but small earthquakes are occasionally recorded. Emerged Quaternary coastal features are found only in the south (Cabinda and Zaïre), where because of slight uplift, some Holocene beaches are present above sea level. This uplift towards the south is also shown by the Miocene coastline, which is at −80 m off Gabon and rises to outcrop on the Cabinda coast near Landana. Radiocarbon dating of brackish clays and freshwater peaty muds uncovered on the shore by wave action indicated a Holocene marine transgression to present sea level about 5,000 years ago, followed by a short regression between 4,000 and 3,000 bp and an oscillation between 3,000 and 500 bp which may have exceeded present sea level by some decimetres, then recent emergence. Mean tide ranges are relatively small (1.0–1.2 m), and slightly higher (up to 1.6 m at Libreville and 1.3 m at Pointe Noire) during spring tides. Tide ranges diminish gradually from north to south along the coast. The climate is tropical, with high rainfall in Gabon (about 2,510 mm per year at Libreville, where mean

monthly temperatures range from 24° to 27°) south to Cape Lopez because of the warm Guinea current from the north. There is much lower rainfall south of Cape Lopez (Banana, at the mouth of the Congo, has 700 mm per year) because of the cool Benguela current from the south, and a dry winter season develops between mid-May and midSeptember. During the summer, warm water from the Guinea current flows towards the south and overlies cold water from the Benguela current, which comes to the surface in winter, when the northerly currents are dominant. The coastal climate is cooler and drier southward through Congo and Babinda to Zaïre in the zone of southeasterly trade winds. Swell from the Southern Ocean has a long period (9–15 s) and a mean amplitude of 1 m (increasing to 3 m during storms), and arrives generally from a southwesterly direction. This swell reworks sands derived from the Serie des Cirques, and has shaped gently curving beaches which are subject to alternations of storm wave erosion and fine weather accretion. The cutting back of some beaches provides sand that drifts generally northward to be deposited in spits and barrier beaches. The beach sand is quartzose with some feldspar and shelly material (less than 5% carbonate). During storms quartz gravels are thrown up to 2 m above high water mark. Beach rock is exposed locally and occasionally (Rémy 1954). The sandy beaches are backed by a succession of beach ridges (2–3 m high) and dunes that extend 20 m above present mean sea level. They include dunes that were built by onshore winds during a pre-Holocene low-sea level phase when the climate was more arid, as well as more recent beach ridges. Older dunes have been podzolised, the upper leached white sands being reworked by wind action about 2,000 years ago. Cliffs and bluffs are of limited extent, but sand-strewn rocky shore platforms occur on headlands. Swampy coasts are extensive, especially on the Gabon coast, but further south they are confined to estuaries and lagoon shores. Mangroves have diminished in recent decades, partly as the result of embanking and land reclamation (Moguedet 1980), but they are spreading in areas of

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.12, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Gabon, Congo, Cabinda and Zaïre

⊡⊡ Fig. 13.12.1 Predominant coastal landforms. (Courtesy Geostudies.)

Gabon, Congo, Cabinda and Zaïre  

⊡⊡ Fig. 13.12.2 Coastal geology. (Courtesy Geostudies.)

   13.12

959

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Gabon, Congo, Cabinda and Zaïre

alluvial deposition, as in the mouth of the Congo, where a rich and varied mangrove forest is developing on the islands and the fringes of the estuary. Most of the sediment delivered by the Congo (40–50 million cubic metres per year) is fine-grained and swept offshore; a little is carried towards the north and is deposited on the continental shelf (Giresse and Kouyoumontzakis 1971). Apart from the formation of shoals and mangrove terraces the Congo sediment contribute very little to coastal sedimentation.

2.  The Gabon Coastline South from the border of Equatorial Guinea the Baie de Corisco coastline is sandy and swampy, with mangroves extensive in the Mondah inlet, sheltered by the Cape Esterias peninsula. Libreville stands on the western side of this peninsula, beside the estuarine bay of Gabon, which is fed by the Mbet and Abango rivers. From Point Pongara the coast southward is again sandy and swampy past Pointe Gnonié to the Baie de Nazare. Here mangroves fringe several inlets in the lee of elongated Cape Lopez. This cape marks the beginning of a drier coastline, fringed by sandy beaches and barriers shaped by southwesterly swell, with beach ridges and backing lagoons and swamps, as at Ogooué and the Tchonga Tchine and Iguéla lagoons. There are local outcrops of phosphatised and silicified Cretaceous limestone. At Setté Cama another lagoon, swampy Lake Ndogo, opens to the sea through a gap in a beach ridge plain. The hinterland remains swampy southeast past the mouth of the river at Nyanga, where the Doungoui River runs parallel to the coastline, behind a sandy barrier, at the southern end of which is Pointe Panga. There are rocky platforms of Pre-Cambrian dolerite at Pointe Kouango, off Mayumba. The town stands at the mouth of the long M’banio Lagoon, which extends behind a 60-km long sandy barrier spit. There are outcrops of ­gabbro on the inner bank. At Pointe Banda Cretaceous (Senonian) limestones outcrop on the shore.

3.  People’s Republic of Congo Across the boundary of the People’s Republic of Congo sandy beaches and dunes fringe the coast to Conkouati, at the mouth of a lagoon fed by the Niambi River, which has built a lagoon delta. The beaches and dunes continue

to the Noumbi River and Pointe Kounda, where there are shore platforms of Cretaceous limestone, generally covered by beach sand. In the hinterland are receding bluffs cut in the Plio-Pleistocene piedmont (Serie des Cirques). A long sandy coastline, backed by beach ridges and dunes, extends southeast past Madingo-Kayes to the mouth of the Kouilou River, which drains through a large mangrove swamp, dominated by Rhizophora spp. and Avicennia spp. At the river mouth are intricate sand spits formed by the interaction of swell and river outflow currents: peaty clay deposits accumulate between and behind these spits. To the south there are much-dissected hills at Diosso, cut into the Plio-Pleistocene Serie des Cirques Formation. There has been beach erosion in recent decades (>Fig. 13.12.3). The beach ends at Pointe Indienne, where Cre-taceous (Turonian) calcareous sandstones and limestones outcrop on the shore. On Pointe Indienne small lagoons are enclosed by sand spits, but these shrink as they are overrun by swash-piled sand. There are small shelly middens dating from the colonial period. The Songololo River cuts through the coastal dune fringe to reach the sea in the next bay (>Fig.  13.12.4). To the south is Pointe Noire, where the Cretaceous (Santonian) dolomites form a shore platform on which oysters grow. Sand bars tend to obstruct the harbour mouth at Point Noire, where continual dredging is necessary. There are archaeological sites. South of Pointe Noire, fishermen periodically open an artificial entrance to a lagoon by bulldozer and thus allow marine fish to enter, but in a few days the barrier is rebuilt. Sandy beaches continue to the mouth of the Loemé River, where the Cabinda border crosses the coastline. Cabinda is an outlying province of Angola, with a coastline about 80 km long. It is fringed by sandy beaches and dunes, interrupted by the mouth of the Massabe lagoon and the Shiloanga River. Near Landana there are shore platforms cut in soft Eocene and Miocene limestones, and to the southeast the sandy beaches curve out past the town of Cabinda to the promontory at Pointe de Tafe.

4.  Zaïre To the south is the Cabinda boundary with Zaïre, now the Democratic Republic of Congo, which has a short ­coastline around the estuary of the Congo River, the mouth of which is partly obstructed by a sand pit that has grown north from the Angola shore.

Gabon, Congo, Cabinda and Zaïre  

   13.12

⊡⊡ Fig. 13.12.3 Beach erosion near the Songololo River mouth, with sand washed up into savanna woodland. (Courtesy G. Kouyoumontzakis.)

⊡⊡ Fig. 13.12.4 Cliff cut by the Songololo River exposing barrier sands. (Courtesy G. Kouyoumontzakis.)

References Giresse P, Kouyoumontzakis G (1971) Geologie du sous-sol de PointeNoire et des fonds sous-marins voisins. Annales Université de Brazzaville 7:97–114 Giresse P, Kouyoumontzakis G (1985) Gabon, Congo, Cabinda, and Zaїre. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 625–638

Moguedet G (1980) La mangrove du Congo. Les Rivages Tropicaux – Mangrove d’Afrique et d’Asie. Talence 1–21 Rémy JM (1954) Contribution a l’étude de la terrasse marine de la Pointe Kudevele (Congo). Rev Zool Bot Africa 48:24–26

961

13.13 Angola

André Guilcher

1.  Introduction Angola has a coastline about 1,600 km long (>Fig. 13.13.1), much of it a narrow coastal plain fronting the Great Escarpment, which is dissected by numerous rivers that descend steep valleys and flow out across the coastal plain to the Atlantic Ocean. The Great Escarpment has been formed by warping of the western margin of the PreCambrian shield that frames Angola (Jessen 1936; Feio 1964), and is bordered by seaward-dipping Cretaceous ⊡⊡ Fig. 13.13.1 The Angola coast. A, location of spits and direction of growth.

and Tertiary formations, which outcrop on the coast. The Tertiary formations include Pliocene red clays and sands near Luanda and Neogene marls and limestones in the surroundings of Benguela and Mocamedes. The seawarddipping sedimentary formations have been truncated by Pliocene and Pleistocene abrasion surfaces, found up to 200 m above sea level and probably uplifted tectonically. Except for the far north, near the mouth of the Congo River, the climate of the Angola coast is semiarid or arid (Guilcher et al. 1974). Mean annual rainfall diminishes rapidly southward to Luanda, which has 323 mm, 200 mm at Lobito, 50 mm at Mocamedes, 20 mm at Porto Alexandre, and 15 mm at Baia dos Tigres and Foz do Cunene. Luanda has a mean monthly temperature of 25.6°C in January and 20.6°C in July, while the equivalent figures for Lobito are 25°C and 20°C. The southerly trade winds blow roughly parallel to the coast and create an upwelling that brings cold water to the surface. This cold water travels to the north as the Benguela coastal current, stabilising the air and forming fog (the cacimbo), but little rainfall: conditions similar to those off Peru and northern Chile, and southern and Baja California. Inland, the rainfall increases rapidly as the influence of marine upwelling fades. Mean spring tide range at Luanda, Benguela, and Mocamedes is 1.2 m. Much of the Angola coast is cliffed and rocky (Guilcher 1985), cut in Cretaceous, Tertiary, and Pleistocene formations, an unusual variety of cliffs and shore platforms in sedimentary rocks in the dry tropics. The beaches are sandy and subject to northward longshore drifting, not because of the trade winds, but because the coast receives powerful southwesterly swell (kalema) originating in the Southern Ocean storms, which breaks obliquely on the shore (Guilcher 1954). This longshore drifting is responsible for the construction of several major spit formations, shaped by southwesterly swell (Guilcher et al. 1974).

2.  The Angola Coastline South from the Zaire border, the coast is low-lying and beach-fringed. Luanda Spit (>Fig. 13.13.2), which shelters

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.13, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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the major seaport of Angola, is decaying. It originally formed before the Palmeirinhas Spit to the south, had grown to its present size. The Luanda Spit extended northeast from Sao Joao de Cazango Island, which is now a dissected remnant within the Palmeirinhas lagoon, bearing a set of successively formed parallel beach ridges that are impossible to explain in their present lagoonal situation. When Palmeirinhas Spit grew to shelter the Luanda Spit, it prevented any further substantial supply of sand. As a result, the Luanda Spit became dissected and cut off from the mainland. Although it is no longer receiving a sand supply, there is still longshore drifting to the north-eastern end, which continues to grow. South of Luanda and behind the Palmeirinhas spit former cliffs cut in the Pliocene red clays and sands are now

bluffs, dissected by numerous active gullies, and separated from the sea by a prograded plain with successively built sandy beach ridges and dunes. Palmeirinhas Spit (>Fig. 13.13.2) is 34 km long. It has grown along the coast northeastward, with the addition of successive ridges parallel to its ocean (NW) shore, and is still growing. It shelters the Palmeirinhas Lagoon, the shores of which have numerous small active spits formed by the small waves generated by local winds. To the south, the River Cuanza spit, 3 km long, is occasionally breached by discharging river floods or strong ocean swell, and the River Longa spit deflects the mouth of the river northward. Between the mouth of the Longa River and Lobito cliffs, some tens of metres high occur on for distances on either side of Three Points Cape.

⊡⊡ Fig. 13.13.2 Luanda Spit and Palmeirinhas Spit, Angola.

Angola

     13.13

⊡⊡ Fig. 13.13.3 Tiger Spit and Bay, southern Angola. Dots: dune ridges, with heights in metres. Dashed line in the south: inner end of the spit tied to the mainland before the 1962 cut. P ON T A D A M A R CA.40 .65

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Lobito Spit, 8 km long, shelters the harbour of Lobito. It was supplied with sandy sediment from the Catumbela River, but this supply diminished after a dam was built across the river, and the delta of is now eroding. Lobito stands on a small cliffed promontory, and at Serra do Sombreiro, near Benguela, there are cliffs 30–50 m high cut in calcareous marls, in which marine corrosion has carved tafoni, notches, visors and pinnacles. Pocket beaches occur between rocky headlands. At Benguela, the coast swings westward to Ponta de San José, then southwest to Ponta das Salinas. The coast is mainly cliffed, with gaps at the mouths of river valleys. Around the Bay of Mocamedes, the cliffs are capped by two sets of Pleistocene beaches, one at 40–50 m, another at 4–5 m (Soares de Carvalho 1961). These emerged beaches have been hardened by calcrete (a calcareous crust) and thus incorporated in the country rock, which is cut by low, jagged cliffs. Near Porto Alexandre, two spits have grown northward: Ponta do Enfiao and Ponta Brava. The latter is 5,500 m long and shelters a fishing harbour. Before it was severed from the mainland in 1962, the large spit enclosing Tiger Bay (Baia dos Tigres) was the longest structure of this kind (37 km) on the coast of Africa. The sand deposits are up to 60 m thick at the northern end, Ponta da Marca, where they form a huge mass on the continental shelf. The sand has drifted along the coast from the south,

across the Namibian border, derived partly from the erosion of desert coasts and partly from the Orange River. Because of the very arid climate, the spit is completely devoid of vegetation, as are the high sand dunes that stand on the landward side of Tiger Bay. The gap that truncated the spit in 1962 will probably not be healed, because the new channel has been deepened by tidal currents, and is bordered by smaller set-back paired spits, shaped by waves generated by the dominant local SSW wind and the southerly trade winds as well as the southwesterly swell. The Tiger Bay spit has thus become a barrier island without a longshore sand supply, and is being re-shaped by the southwesterly swell. On the landward side of Tiger Bay, the coast has a number of small hooked spits directed northward (>Fig. 13.13.3). The hinterland is a desert with numerous NW-SE sand ridges 30–50 m high, formed by southerly and southwesterly winds. Some 30 km to the north, these dune ridges take the shape of giant waves divided into barchans moving northwestward (>Fig. 13.13.4). Cunene River, which flows along the border between Angola and Namibia, is a perennial stream fed by tropical rainfall in the hinterland and flowing down to an arid coast. The mouth of the river (Foz do Cunene) has been deflected to the north by the incident SW swell. Former parallel courses of the river can be seen a short distance inland. ⊡⊡ Fig. 13.13.4 Large sand waves containing multiple barchans, northeast of Tiger Bay. They are truncated along the coast by a wide beach and dune ridge, which has the appearance of a foredune, but bears little or no vegetation.

Angola

References Feio M (1964) On the evolution of the great escarpment in the southwest of Angola (in Portuguese). Garcia de Orta 12:323–354 Guilcher A (1954) Dynamics and morphology of the sandy coasts of Atlantic Africa (in French). Cahiers d’Information Géographique 1:57–68

     13.13

Guilcher A (1985) Angola. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 639–643 Guilcher A, Medeiros CA, Matos JE, Oliveira JT (1974) The restingas (longshore spits) of Angola (in French). Finisterra 9:117–211 Jessen O (1936) Travels and research in Angola (in German). Reimer, Andrews and Steiner, Berlin Soares de Carvalho G (1961) Geology of the Mocamedes desert (in Portuguese), vol 26. Junta Investigacione Ultramar, Memoir

967

13.14  Namibia

Eric Bird · A. Goudie · H. Viles

1.  Introduction The Namibian coastline is about 1,572 km long, extending from the Kunene River in the north (17°16'S) to the Orange River in the south (28°33'S) (Bremner 1985). Except for a broad embayment at Walvis Bay, the coastline is fairly straight, with a NNW trend. Two natural harbours capable of handling medium-tonnage vessels are located at Walvis Bay and at Luderitz. The former is protected by Pelican Point, a large sand spit, and the latter is situated in a NW-facing rocky inlet. The Namibian coast, a passive margin feature that dates back to the early Cretaceous (Goudie and Eckardt 1999), consists predominantly of late Pre-Cambrian Da­­ mara metasediments (green schist to amphibolite facies), a geosynclinal sequence with associated granitic intrusive rocks. The foliation of the Damara metasediments is roughly parallel to the coast between the Kunene and the Ugab rivers, but southward to Meob Bay it swings to strike at right angles to the coast. A number of granitic intrusions along the core of the Damara geosyncline outcrop along the coast, as at the mouth of Messum River and on  Cape Cross. Palaeozoic rocks are restricted to the Dwyka glacial beds that fill a valley in the Pre-Cambrian basement near Cape Fria, and Mesozoic basic lavas occur sporadically between Cape Fria and in the vicinity of Palgrave Point. Mesozoic marine sediment are represent­ ­ed by a single outcrop of sandy limestone near Bogenfels, and there are several remnants of Tertiary marine deposits. Much of the coastal fringe is blanketed by unconsolidated Quaternary sand, particularly adjacent to the Namib Sand Sea. Beach sediment are predominantly medium to finegrained sand, consisting of quartz, feldspar, heavy minerals and some calcium carbonate. The proportion of heavy minerals, such as garnet and ilmenite, increases substantially (up to 80%) near rocky foreshores, from which they are presumably derived. The Orange River is a major source of sediment supply (mainly sand and mud, including diamonds) to the coast: its sediment load is said to be five times that of the River Nile (Lancaster and Ollier 1983).

The beaches and coastal dunes of light brown sand originate from sediment supplied by the Orange River, and contrast markedly with the semi-stable reddish dunes of the interior. Strong northward longshore drifting is indicated by several spits along the coast, and the occurrence of diamonds in decreasing quantities north of the Orange River mouth. Northward drifting results from the arrival of a southwesterly ocean swell (wave period typically 6–9 s) and southerly waves generated by trade winds (wave period typically 4 to 7 s). Ocean upwelling associated with the cool Benguela current results in a mild, dry, windy and foggy coastal climate. Walvis Bay has a mean temperature of 14°C in August, rising to 19°C in February, and a mean annual rainfall of about 25 mm. Fog occurs on 100 days in the year around Swakopmund and precipitates appreciable amounts of moisture. Strong berg winds from the east blow plumes of dust into the Atlantic (Eckardt et al. 2001). Tides are semidiurnal with a mean spring tide range of 1.4 m at Walvis Bay, but co-tidal lines run almost parallel to the coastline and tidal currents are weak, producing little sediment transport.

2.  The Namibian Coastline From the Angolan border on the Kunene River to Walvis Bay the Skeleton Coast (>Fig.  13.14.1) is backed by migrating sand dunes (Lancaster 1982), salt pans and deflation surfaces. The coast consists of sand or gravel beaches strewn with whale bones and driftwood (which on this arid coast must have drifted from forested regions far north) and interspersed with rocky headlands. In places the coastline may be as much as 2 km seaward of the charted position because of progradation. The Skeleton Coast Park was established in 1963. The estuary of the perennial Kunene River is broad and shallow, with a sand bar across its mouth in the dry season. It cuts through the Kunene Sand Sea, which extends northward into Angola. From the Kunene River to Cape Fria the coastal waters are strewn with imperfectly charted shoals and rocky reefs of Damara metasediments.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_13.14, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 13.14.1 The Skeleton coast of Namibia showing the Sand Sea directly adjacent to the Atlantic.

⊡⊡ Fig. 13.14.2 The mouth of the Swakop River at Swakopmund showing a railway bridge that was destroyed by a flood in the 1930s.

The beach is mainly sandy and the backshore low-lying, with large, elongated coast-parallel salt pans. At Cape Fria low hills of Cretaceous basalt end in cliffs and ledges of lava behind the beach. From Cape Fria to Möwe Point the shore is sandy, with fringing rocky reefs at False Cape Fria. Rocky Point is a prominent headland of the basalt that projects some 600 m

into the sea, and is backed by terrain which is flat and sandy, with a few small salt pans. Möwe Point is a low foreland of Damara metasediments, and heavy surf prevails along the coastline to the south. From Möwe Point to Palgrave Point the coast is also low-lying, with scattered salt pans, small shifting sand dunes and scrubby vegetation. The shore is rocky,

Namibia

c­ on­sisting of Damara metasediments overlain by, and often fault­­ed against, Cretaceous basalt. Rivers such as the Hoanib and the Unjab, which opens to Torra Bay, flow only rarely. They are frequently blocked by banks and coastal dunes, leading to ponding and silt deposition. From Palgrave Point to Cape Cross the Damarland coast has low cliffs separated by long sandy beaches. Numerous salt pans lie in the backshore. Damara metasediments alternate with massive Cretaceous basalt. The foreland at Cape Cross projects some 5 km into the sea and has low rocky cliffs and a nearby reef. A cross was erected here by the Portuguese navigator Diego Cão in 1486, but the existing cross is a recent replica. The cape is the site of a major seal colony. The coast from Cape Cross to a point roughly 50 km north of Walvis Bay is generally rocky, with numerous shoals close inshore. The rocks consist mostly of Damara metasediments with occasional granitic intrusions. The lower reaches of the Omaruru River cross an extensive cover of Quaternary sand that extends northeast along the Damara foliation trend (Miller and Schalk 1980) before entering the resort of Henties Bay. The coast is low-lying, with a few small salt pans. South of Cape Farilhao is the sand-choked harbour at Swakopmund, the northern limit of diamonds in the coastal sand deposits that have drifted from the Orange River mouth. The Swakop River (>Fig. 13.14.2), which enters the Atlantic at Swakopmund, is prone to flooding and as a result of a flood event in 1934, prograded the coastline by over 1 km. Walvis Bay was formerly a whaling port. The sand spit protecting it is about 18 km long and grew northward by about 1,800 m between 1905 and 1985. It has a maximum elevation of about 13 m, and terminates at Pelican Point. Large mud islands occasionally appear in the bay off Pelican Point because of the extrusion of soft diatomaceous mud from beneath the accumulating sand load. Hydrogen sulphide released from the mud has caused periodic mass mortality of fish in this region. There are extensive marshlands near Walvis Bay, which has saltworts. Platforms have been built offshore for seabirds to roost and nest, and thus deposit large quantities of guano, which are periodically harvested for use as agricultural fertiliser. The Kuiseb River forms the northern boundary of the Namib Sand Sea. It reached the sea only three times in the twentieth century (1933, 1962/63 and 1985) but welcomed the new millennium by flowing to the coast in 2000. The Namib Sand Sea (>Fig. 13.14.3), an area of sandy desert (erg) with patches of gravel (hamada), backs a desolate, forbidding coast between Walvis Bay and Luderitz, a distance of about 400 km. It extends inland for about 100 km to the Great Escarpment. The coast consists partly of

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⊡⊡ Fig. 13.14.3 Northern Namib Sand Sea with forelands and spits. (Courtesy J.M. Bremner.)

l­ow-lying, sparsely vegetated sand flats and partly of dunes up to 150 m high. This is the Sperrgebiet. The northern half of this section has four large depositional formations, at Walvis Bay, Sandwich Harbour, Conception Bay and Meob Bay. These are situated adjacent to each other to former river estuaries: it is thought that the Swakop River once entered the sea at Walvis Bay, the Kuiseb River at Sand­­wich Harbour, while the two vleis, Tsondab and Sossus, breached

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the intervening dune barrier to reach the sea at Conception and Meob bays respectively (Bremner 1985). Northward drifting sand has progressively blocked the four former estuaries, the sequence being illustrated by a decrease in salt pans and an increase in water areas from south to north. The depositional forelands are essentially spits, with low sandy bluffs along their outer coasts, with estuarine areas behind. In between are high walls of sand, as at Die Lange Wand. Parts of the spits are prograding, as  shown by the wreck of the Eduard Bohlen, which grounded in 1910 and now lies roughly 400 m inland from the beach. Between Conception Bay and Meob Bay, metasediments  of  the Damara Sequence and granitic intrusions outcrop sporadically between the dunes and on the shore. The modern erg (sandy desert) is underlain by a lithified erg, the Tsondab Sandstone, that dates back to the early Tertiary. From St. Francis Bay to Luderitz, the coast of the Namib Sand Sea is more indented, with numerous small rocky headlands and intervening sandy bays. There are five small islands close inshore, and rocky reefs commonly surround them. North of Spencer Bay, the rocks comprise Damara metasediments, with a coast-parallel foliation trend, whereas southward they are basement gneisses of pre-Damara age, about 1,200 million years old. South of Hottentot Point, about 40 km NNW of Luderitz, is the rocky island of Ichaboe, which in 1843 had a capping of guano up to 7.5 m deep, deposited by penguins and gannets. Between 1843 and 1850 about 800,000 tonnes of

guano were removed and exported as fertiliser, exposing the rocky basement. Luderitz lies in a north-facing rocky inlet containing several small guano-plastered islands, notably Flamingo Island, Seal Island, Penguin Island, Shark Island and Halifax Island. The rocks are predominantly basement gneisses and schists of pre-Damara age. The town became prosperous with the discovery of diamonds in 1908. Elizabeth Bay, 30 km south of Luderitz, is flanked on either side by a rocky coast with submerged reefs of preDamara acid/basic intrusive rocks. Southward of Chamais Bay, the coast continues to be indented with small rocky headlands and intervening sandy bays, and the coastal waters are dotted with seven small guano-capped islands. At Bogenfels, midway between Lüderitz and Chamais Bay,  dolomites of the Gariep Complex (equivalent in age to the Damara Sequence) have been eroded into spectacular cliffs and arches (>Fig. 13.14.4). The only marine Cretaceous sedimentary rocks to be found along the west coast of southern Africa are in the Sperrgebiet at an altitude of 60–70 m (Haughton 1963). In the same area, a few small exposures of Eocene marine sediment lie between 70 and 140 m above sea level. From Chamais Bay to the mouth of the Orange River, the coastline is straight, with a backshore covered with low dunes. The region is intensively mined for alluvial ­diamonds, and a sizable town, Oranjemund, located 6 km north of the river, has been established to support the industry. The Orange River is about 1.5 km wide at its

⊡⊡ Fig. 13.14.4 The Bogenfels Arch, 51 m high, has been cut through a recumbent syncline of dolomite belonging to the late Pre-Cambrian Gariep complex. (Courtesy J. Rogers.)

Namibia

mouth, but during the dry season it becomes completely blocked by a wave-built sand bank. Emerged beaches near the Orange River at an elevation of 25 m have yielded warm-water Pleistocene molluscs. The South African boundary runs along the Orange River.

References Bremner JM (1985) Namibia. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 645–651 Eckardt FD, Drake N, Goudie AS, White K, Viles H (2001) The role of playas in pedogenic gypsum crust formation in the Central

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Namib Desert: a theoretical model. Earth Surf Process Landf 26:1177–1193 Goudie AS, Eckardt F (1999) The evolution of the morphological ­framework of the Central Namib Desert, Namibia, since the early Cretaceous. Geogr Ann 81A:443–458 Haughton H (1963) Stratigraphic history of Africa South of the Sahara. Oliver & Boyd Ltd., London Lancaster N (1982) Dunes on the Skeleton Coast, Namibia: geomor­ phology and grain size relationships. Earth Surf Process Landf 7:575–587 Lancaster N, Ollier CD (1983) Sources of sand for the Namib Sand Sea. Z Geomorphol Supplementband 45:71–83

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14.0 South Africa – Introduction Gerald G. Garland   

The South African coast is some 3,000 km long, fronting both the south Atlantic and South Indian Oceans. About 1,700 km of the coastline consists of sandy beaches, the rest being either rocky or a mixture of rocks and sand. Two-thirds of the estuaries are on the east coast between Cape Padrone and Mtunzini. Numerous reports have been published by CSIR (Stellenbosch) and others describing individual estuary and lagoon systems. On the east coast, especially, sediment yields from the main rivers are high, sustaining sandy beaches and barriers. The arid west has far fewer rivers and consequent low sediment delivery. Only 19 offshore islands occur, mainly in the west and off Port Elizabeth. The country is too far south for true coral reefs to develop, although the coral-clad structures north of Sodwana in the east exhibit many characteristics of tropical reefs. Mangroves, although not extensive, are found as far south as East London (33°S). Coastal dunes are well distributed along much of the coastline. Apart from small dunes common on rocky coasts and the classic parabolic dunes of the south and south-east, there are high, bare, and mobile dunes on the west coast, and stable, weathered, and vegetated dunes up to 80 m high on the east coast, continuing into Mozambique. These are thought to be of late Pleistocene or early Holocene age. Subsequent to the break-up of Gondwana, South Africa’s coastal position had stabilised by the end of the Cretaceous, so coastal landform assemblages tend to reflect the nature of underlying lithology, sea level fluctuations in the Tertiary and Quaternary, and modern ­processes. There are sectors of cliffed and rocky coast, sometimes fronted by narrow beaches of sand or gravel, on the west and successions of headlands and bays, with beaches backed by dunes, several river mouths and lagoons, on the east. There are indications of several relict coastlines in South Africa. Although dating is difficult, strandlines on  the west coast have been observed at 75–90, 45–50, 29–34, 17–21, 7–8, 5, and 2 m above present datum. On the east coast emerged (raised) beaches that are present at 60  and  30 m above present sea level, with relict

coastlines  at 8, 3.4–5.3, and 4.5 m on the shores of the coastal lagoon, Lake St. Lucia. A sea level 68 m lower than present is indicated by submerged dune ridges off the east and south coasts. Recent investigations of late Quaternary (Ramsay and Cooper 2002) and Holocene (Miller et al. 1995; Ramsay 1996) sea level have revealed major changes broadly consistent with global eustatic trends. During the late Holocene, however, geomorphological and sedimentological evidence points to a series of minor fluctuations around the present level since about 7,000 years bp. The best data on contemporary rates of sea level change show a mean annual rise in sea level between 1.7 mm (west coast) to about 3.0 mm (east coast) (>Fig. 14.0.1). Climatic variation along the South African coast is largely attributable to the major oceanic currents (>Fig.  14.0.1). The difference between the west, controlled by the northward-flowing Benguela current bringing cold waters from the South Atlantic, and the east, influenced by tropical waters carried south by the Agulhas current, is exemplified by a climatic comparison between Port Nolloth in the west and Durban in the east. Although within a few minutes of latitude of each other, Port Nolloth, with a sea surface temperature rarely exceeding 16°C, has a mean annual rainfall of 45 mm and a mean annual temperature of 14°C. Comparable statistics for Durban, where sea surface temperatures vary between 20 and 30°C, are 1,018 mm and 22°C respectively, the mean monthly temperatures being 16.7°C in January and 23.9°C in July. The net result is that the west coast from the Namibian border south to Lamberts Bay has a desert climate, the section between Lamberts Bay and Mossel bay on the south coast is broadly Mediterranean and a temperate coastal climate occurs between Mossel Bay and Port Elizabeth, from which the climate is subtropical up the east coast to the Mozambique border. With a mean spring tide range of less than 2 m, the South African coast is microtidal: Saldanha and Cape Town have a range of 1.5 m, Knysna 1.6 m, Port Elizabeth 1.5 m, East London 1.6 m, Durban and Richards Bay 1.8  m. The regular coastal topography with few inlets, straits, or narrows of significant dimensions means that

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_14.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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South Africa – Introduction

⊡⊡ Fig. 14.0.1 South Africa – ocean currents. (Courtesy Allan Heydorn.)

tidal ­channelling is almost entirely absent. Tidal currents therefore have played a very restricted role in geomorphological development. Waves on the other hand are crucial. Ocean swell is predominantly from southwest to southeast throughout the year, a southerly fetch of thousands of kilometres resulting in a moderate to high wave energy coastline. Wave heights are typically between 1 and 3 m with a tendency to decrease northwards, such that the wave height exceeded 10% of the time falls from 10 m off the Cape Peninsula to 4 m at Walvis Bay on the west coast across the Namibian border and 6 m at Richards Bay on the east coast. The wave climate of the South African coast is profoundly influenced by seasonal shifts (southward in summer and northward in winter) of the gale zone of the Antarctic Circumpolar Current region (>Fig. 14.0.1, G), and by local winds generated by the regular passage of atmospheric lows over the coast from west to east. These wave patterns strongly influence coastal processes such as longshore drifting of beach sediment, and have shaped coastal features such as headlands and bays, sand spits and the configuration of river mouths. To the west of Cape Agulhas, the coastline borders the South Atlantic, and to the east the Indian Ocean. The east coast waters are characterised by the warm waters of the

southward flowing Agulhas Current, those of the west coast by sporadic upwelling of cold, nutrient-rich waters typical of the Benguela Current regime. Along the southwest and south coasts, extensive mixing of water masses occurs. Although the main ocean currents influence coastal climate, they have a limited effect on coastal landforms and processes. Instead the dominant swell, combined with local shore currents (especially retroflections of the Agulhas current in the east) as well as regional bearing of the coastline with respect to swell, result in a prevailing northerly drift. Coastal sediment tends to move northward along both the east and the west coasts unless disturbed by local conditions, and between 1 and 2 million tonnes of sediment annually may pass any given point. These powerful drifts and high sediment loads combine to shape the spits and bars, which restrict, and sometimes seal, the mouths of many of east and south coast estuaries. Onshore–offshore sediment motion is also controlled by wave-induced currents. With the exception of the south and south-west coasts where, due to climatic reasons, the situation is reversed, wave-driven asymmetric flux tends to build beaches during the relatively calm southern hemisphere winters and erode them in stormier summer weather. Powerful rip currents, which maintain the mass

South Africa – Introduction

balance of water in the nearshore zone, develop off the long straight beaches of the northeast coast. These may sweep up or down the coast depending on the direction of wave approach, and are responsible for withdrawing sediment beyond the swash zone. The following chapter describes South African coastal landforms in sequence from the Olifants River round to the Mozambique border.

14.0

References Miller DE, Yates RJ, Jerardino A, Parkington JE (1995) Late Holocene coastal change in the southwestern Cape. S Afr Quat Int 29–30:3–10 Ramsay PJ (1996) 9000 Years of sea-level change along the southern African coastline. Quat Int 31:71–75 Ramsay PJ, Cooper JAG (2002) Late Quaternary sea level changes in South Africa. Quat Res 57:82–90

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14.1  South Africa Andrew Cooper

1.  Introduction

northward-directed longshore drift. The river-dominated estuary of the Orange River (Cooper 2001) has a braided channel that extends to a coarse sand and gravel barrier at the river mouth. Most of the bed load of the Orange River passes to the beaches of Namibia to the north. To the south the coastline has rocky and irregular cliffs, subject to substantial wave erosion. Some cliffs plunge directly into the sea; others have narrow sand or gravel beaches along their bases. Beyond the Olifants River the coastline gradually changes from steep rocky shores to increasingly wider-spaced rocky headlands separated by sweeping sand beaches backed by dunes. The metamorphic bedrock is replaced by quartz sandstones belonging to the Table Mountain Group, and this change in bedrock lithology is seen in the composition of local beaches, which consist predominantly of shelly quartz sand devoid of heavy minerals. Port Owen at the mouth of Berg River has parallel breakwaters beside a cut through a barrier enclosing a lagoon. A dramatic change in coastal morphology takes place to the south of the Berg River, where the quartz

The landforms of the South African coast (Heydorn and Flemming 1985) can be broadly subdivided into six physiographic regions: Orange River to Olifants River; Olifants River to Berg River; Berg River to Cape Agulhas; Cape Agulhas to Cape Padrone; Cape Padrone to Mtunzini; and Mtunzini to Ponta do Ouro (>Fig. 14.1.1).

2.  The Coastline South of the Orange River mouth at Alexander Bay the coast consists of low rounded cliffs and rocky shores on Pre-Cambrian metasediments, fronted by narrow beaches of sand or gravel. On the 407 km coastline between the Orange and Olifants Rivers nine smaller rivers enter the sea. In the present climate they flow only rarely, but there are occasional floods. The west coast is exposed to the continuous pounding of heavy southwesterly swell, which produces a strong

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sandstones give place to granites, which dominate the coast between the Berg River and the Cape Peninsula (south of Cape Town). The coast is very irregular and its northern half is characterised by numerous pocket bays, headlands and low rocky islands. The largest embayment, Saldanha Bay, is a natural harbour and has a southern extension known as Langebaan Lagoon, separated from the sea by a barrier of dunes and Pleistocene dune calcarenite, and having shores bordered by salt marsh. The coastline between Saldanha Bay and Table Bay has widely spaced rocky headlands separated by long sandy beaches backed by dunes. Table Bay is bordered by a promontory which marks the northern limit of the Cape Peninsula. The cliffs are dramatic, with very steep slopes cut in layered sandstone above the granite basement, dropping several 100 m and plunging into the sea. The abundance of cliffs and high wave energy has enabled the development of several small gravel and boulder pocket beaches around the coast. The Cape Peninsula is separated from the high coastal ranges of the hinterland by a lowland known as the Cape Flats. It consists of Tertiary and Quaternary sands, which is the source of the beaches along the northern shore of False Bay. The east coast of False Bay is steep (>Fig. 14.1.2), running out to Cape Hangklip, where the coastline swings to the east, initially flanked by towering sandstone ridges. The irregular rocky shore is interrupted by long curving sandy beaches, as at Hermanus, which in many places

form the seaward margins of barred estuaries. The relief gradually diminishes toward Cape Agulhas, although the shore remains predominantly rocky, interspersed with short sandy beaches. Cape Agulhas (>Fig. 14.1.3) marks the southern tip of Africa. To the east the coastal morphology changes markedly, and is dominated by the regional strike of the Cape Fold Belt, in which an alternating sequence of anticlines and synclines is cut at an oblique angle by the general trend of the coastline. As a result, there is a succession of half-heart or crenulate bays opening to the southeast, each with an anticlinal headland followed by an asymmetrical bay occupying a synclinal depression. The crenulate shape of each bay is determined by the prevailing southwesterly swell, refracted round the headlands. To the east of the Gouritz River the character of the coastline again changes, and becomes dominated by a raised coastal platform at a height of 150–250 m, probably late Cretaceous to early Tertiary in age, fronted by a steep coastal slope. The broad beaches locally have barchan dunes migrating alongshore, as at the Bree mouth. There are coastal outcrops of dune calcarenite at Swartvlei. Rivers along this part of the coast are deeply incised. At Knysna the erosion of softer layers has led to the ­formation of a large estuary connected to the sea through a ­narrow rocky channel, while the Wilderness Lake System has developed between successive coastal dunes ­occupying gaps in bedrock ridges. The lakes have been ­segment­­ed 

⊡⊡ Fig. 14.1.2 The steep Berg coast, south of Strand, looking towards Cape Hangklip. (Courtesy Geostudies.)

South Africa

14.1

⊡⊡ Fig. 14.1.3 The rocky shore at Cape Agulhas. (Courtesy Geostudies.)

⊡⊡ Fig. 14.1.4 Plettenberg Bay lagoon and dune-capped barrier. (Courtesy Geostudies.)

into a chain of lagoons, Eilandvlei, Langevlei and Rondevlei, shrinking as the result of reed and rush encroachment. Beyond Cape Seal is the large Plettenberg Bay tidal lagoon, behind a sandy coastal barrier with a central gap which has migrated westward (>Fig. 14.1.4). To the east is a steep coast at Stormsrivier (>Fig. 14.1.5). The city of Port Elizabeth is situated on the coast of Algoa Bay near the northeastern limit of the Cape Fold Belt. Algoa Bay is composed of asymmetrical bay beaches extending to Cape Padrone. The topography immediately

to the west and east of Port Elizabeth is much flatter than that of most of the south coast. The quartzites and shales of the Table Mountain Group are replaced by Cretaceous conglomerates, sandstones, shales and limestones, and by unconsolidated Tertiary and Quaternary sediment. Dunes have spilled northward over Cape Recife, where shore platforms have been cut into Pleistocene dune calcarenite at Flat Rock, and beach rock is exposed after storms on the shore at Summerstrand. The breakwater at Port Elizabeth harbour has intercepted northward drifting sand,

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⊡⊡ Fig. 14.1.5 Steep coast at Stormsrivier. (Courtesy Geostudies.)

so that beaches downdrift have been eroded and concrete tetrapods have been dumped to stabilise the coastline. To the east dunes fringe the coast of Algoa Bay (>Fig.  14.1.6). Large coastal dune fields form backshore zones of unvegetated sand ridges orthogonal to the coastline, moving to and fro in response to seasonally alternating westerly and easterly winds (>Fig. 14.1.7). At Sundays River dunes are spilling into the estuary from the south (>Fig. 14.1.8). At Cape Padrone the coastline swings to the northeast. There is no coastal plain, the coast being dominated by more uniform, convex slopes, except around Port St. Johns, where complex faulting has produced high cliffs and steep slopes. The rocky coast is interrupted by short sandy beaches in the vicinity of river mouths, most of which are encumbered by spits and inwashed sand thresholds. The abundance of estuaries is particularly high between East London and Durban because of the steep coastal hinterland (Cooper et al. 1999; Cooper 2001); many are enclosed by a barrier for long periods, only opening in response to river discharge or overwashing. Most estuaries in this region have mangroves, which are absent on the south and west coasts, where salt marshes including Spartina capensis (>Fig. 14.1.9) are found. Rock outcrops gradually diminish northward from East London until the coast eventually becomes dominated by sandy beaches, river mouths, lagoons and estuaries. Hole in the Wall is a natural arch cut through an elongated stack of horizontally stratified sandstone near

the mouth of a small gravelly river. At Amanzimtoti, some 25 km south of the city of Durban, the sporadically occurring coastal dunes merge into a quasi-continuous ridge, which extends northward beyond the Mocambique border. This ridge contains a core of Pleistocene dune sand. At Durban the beaches are interrupted by The Bluff, a bold headland fringed by shore platforms cut in dune calcarenite. Beach rock has emerged on these shore platforms. The Durban Harbour breakwaters, completed in 1952, have intercepted northward-drifting sand, depriving the beaches to the north of their natural sediment supply. In 1953, Durban City Council began to pump sand northward under the harbour entrance to replenish these beaches, and they have been augmented with sand dredged from the harbour (Laubscher et al. 1990). On the Natal coast north of Durban unconsolidated Quaternary deposits overlie Karoo sediment to form beaches and dunes that are punctuated by occasional, small outcrops of Pleistocene aeolianite (dune calcarenite) or intrusive PostKaroo ­volcanics, as well as by river mouths and lagoons. Extensive drifting sands impede, and occasionally close, river mouths and lagoon outlets. Beach rock outcrops are encrusted below mid-tide level with algae and zooanthids (Miller and Mason 1994). North of the Tugela River the accumulation of beach ridges topped by foredunes has created a rare advancing shoreline (Cooper et al. 1999). The introduction of sand for shoreline advance is closely linked to episodic floods in

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⊡⊡ Fig. 14.1.6 The dune fringe at Algoa Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 14.1.7 Dune crests move eastward and westward with alternating ­coastwise winds. (Courtesy Geostudies.)

the Tugela River which periodically supply large quantities of sand to the coast. Between Mtunzini and Ponta do Ouro is the Zululand Coastal Plain, on which seaward-dipping Cretaceous and Lower Tertiary strata are unconformably overlain by Upper Tertiary to Holocene sediment. The Quaternary history of the area is particularly complex, and field evidence ­suggests

that there have been several superimposed and reactivated barrier lagoon systems in the wake of successive Pleistocene and Holocene sea level fluctuations. At present, the region is characterised by a number of large ­estuaries and coastal lagoons (Wright et al. 2000), which appear to have developed from more extensive ­mid-­Holocene estuaries by ­progressive siltation. It is probable that the evolution of

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⊡⊡ Fig. 14.1.8 Dunes spilling into the estuary at Sundays River. (Courtesy Geostudies.)

⊡⊡ Fig. 14.1.9 Spartina capensis in the estuary of Bushman’s River. (Courtesy Geostudies.)

these estuaries and lagoons has been influenced by a minor Late Holocene fall in sea level. At Richards Bay the northern section is now the harbour and the southern Mhlatuzi lagoon is a sanctuary that has been given a separate entrance. Heavy siltation has occurred because the filtering swamps of the flood plain of the Mhlatuzi River were canalised to facilitate the construction of a railway line. Interception of northward ­drifting sand by the harbour breakwater has resulted in

erosion of beaches to the north (>Fig. 14.1.10), countered by piping sand past the entrance. St. Lucia is a major subtropical coastal lagoon, with bordering mangroves. The artificial stabilisation of the inlet had a profound influence on the dynamics and ecology of this lagoon, but recently, the KwaZulu-Natal Wildlife Service changed its policy and allowed the mouth to fluctuate naturally. As a consequence, the inlet has been closed since 2002. The enclosing sand barrier is up to 5 km

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⊡⊡ Fig. 14.1.10 Cliffed dunes indicate erosion behind the depleted beach north of Richards Bay. (Courtesy Geostudies.)

wide with beach rock outcrops that are salient at Cape Vidal (Cooper 1991). Beach rock outcrops are found intermittently along the sandy shoreline north of St Lucia. Larger outcrops of beach rock and/or dune calcarenite form headlands that have enabled the formation of log spiral bays on their downdrift side. Long sandy beaches backed by high vegetated dunes extend between St. Lucia and the Mozambique border. The only breaks are where occasional outcrops of dune calcarenite or beach rock are present. A large barrier encloses Lake Sibaya, formerly an estuary and now a freshwater lake (Wright et al. 2000). Kosi Lagoon, close to the Mozambique border, is separated from the sea by a high, vegetated, dunecapped barrier, with a broad sandy beach on its seaward side at Bhanga Nek. The lagoon has been segmented into a series of linked basins of varying salinity and circulation patterns. The tidal inlet through which the lagoon exchanges water with the ocean has been recorded to close only once, after a tropical cyclone, and it was artificially reopened. Coral reefs are present offshore (Ramsay and Mason 1990) as patch reefs on submerged dune calcarenite and beach rock ridges.

References Cooper JAG (1991) Beach rock formation in low latitudes: implications for coastal evolutionary models. Mar Geol 98:145–154 Cooper JAG (2001) Geomorphological variability among microtidal estuaries from the wave-dominated South African coast. Geomorphology 4:99–122 Cooper JAG, Wright CI, Mason TR (1999) Geomorphology and sedimentology of South African estuaries. In: Allanson BR, Baird D (eds) Estuaries of Southern Africa. Cambridge University Press, England, pp 5–25 Heydorn AEF, Flemming BW (1985) South Africa. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 653–667 Laubscher W, Swart DH, Schoonees D, Pfaff WM, Davis AB (1990) The Durban beach restoration scheme after 30 years. In: Proceedings 22nd coastal engineering conference, vol. 3, Delft, pp 3227–3238 Miller WR, Mason TR (1994) Erosional features of coastal beach rock and aeolianite outcrops in Natal and Zululand, South Africa. J Coast Res 10:374–394 Ramsay PJ, Mason TR (1990) Development of a type zoning model for Zululand coral reefs, Sodwana Bay, South Africa. J Coast Res 6:829–852 Wright CI, Miller WR, Cooper JAG (2000) The Cenozoic evolution of coastal water bodies in northern KwaZulu-Natal, South Africa. Mar Geol 167:207–230

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15.0 East Africa – Editorial Introduction

Introduction The East African coast, facing the Indian Ocean, is generally low-lying, with coastal plains over basins of Mesozoic and Tertiary sedimentary rock bordering uplands of ­Pre-Cambrian Shield formations (Orme 2005). A zone of N-S  faulting oc­­curs at the southern end of the East African  Rift, passing beneath the Zambezi delta. The > Mozambique coast is exposed to SE ocean swell from the Indian Ocean and waves generated by SE and NE winds. Mean spring tide ranges are between 3.5 and 4.5 m, increasing to 5.6 m at Beira. Northern Mozambique, in the lee of Madagascar and thus sheltered from Indian Ocean swell, has cliffs cut in Cretaceous and Tertiary rocks on an embayed coast, fringed by coral reefs, some emerged. A coastal lowland of varying width extends northward into Tanzania. The coasts of > Tanzania and > Kenya have warped and faulted marine terraces resulting from Quaternary tectonic movements. There are beaches and dunes, barriers, lagoons and mangrove swamps, and coral reefs, some of which have been uplifted to form coastal terraces. Then islands of Mafia, Zanzibar, and Pemba are emerged coralline formations. In northern Kenya and > Somalia, increasing aridity is marked by a widening desert hinterland. Quaternary dunes ­border the coast on

either side of Mogadishu, and north of Obbia cliffs up to  90 m high are cut in Tertiary limestones. Southeasterly swell from the Indian Ocean is accompanied by waves produced by easterly to southerly winds, except in the winter NE monsoon. Mean spring tides are about 2 m. Cliffs alternate with alluvial plains along the Gulf of Aden coast in Northern Somalia, where marine terraces and emerged coral reefs indicate Quaternary tectonic deformation. As the coast curves northward to Djibouti, it becomes more arid. On the west coast of the Red Sea, the Pre-Cambrian basement comes close to the coast in a dissected mountain front, fringed by a narrow coastal plain in > Eritrea and the > Sudan. Coral reefs border the coast, and there are emerged coral terraces. Mean spring tide diminishes to less than 0.5 m and wave action is determined by NE trade winds over the Red Sea. It is convenient to include the Sinai Peninsula here, along with the short sectors of Israeli and Jordanian coast at the head of the Gulf of ‘Aqaba.

Reference Orme AR (2005) Africa, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 16–21

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

15.1 Mozambique

Maria Eugénia Soares de Albergaria Moreira

1.  Introduction The Republic of Mozambique on the east coast of Africa has a coastline more than 2,470 km long, more than 90% of which is a low coastal plain of sand or mud stabilised by plants (Tinley 1985). The lowlands are up to 200 km wide in the southern part of the country, narrowing as up­­ lands approach the coast in the northern part. Triassic and Cretaceous formations dip seaward over a Pre-Cambrian basement (of metamorphic gneisses, migmatites, and am­­ phibolites), which outcrops locally on the coast. The southern coast has beaches backed by several generations of dunes and lagoons. Some of the dunes have become consolidated as dune calcarenites (aeolianites) that outcrop in headlands. Generally, the dunes are cliffed with blowouts on the seaward side, and forested parabolic dunes dominate the inland topography. Beach rock outcrops on the beaches, and there are submerged beaches and dunes marking Quaternary coastlines. The central coast has headlands between arcuate bays with beaches and barriers estuarine lagoons and the Zambezi delta. The northern coast is linear, with cliffs and bays cut out along fault lines in emerged coral. Along the narrow shelf, a fringing reef emerges as a coral strand with small coralline islands that form the Quirimbas archipelago. The climate is tropical, with hot, wet summers (December–March) and drier winters. The southern coast has southeasterly trade winds, occasional heavy rain and drought periods, but rainfall increases northward and the coast has a monsoonal regime and occasional tropical cyclones (hurricanes). Maputo has a mean monthly temperature of 18.3°C in July, rising to 25.6°C in January and an average annual rainfall of 760 mm, while Beira, to the north, has 20.6°C in July, 27.8°C in January and an average annual rainfall of 1,500 mm. The country lies due west of Madagascar, from which it is separated by the Mozambique Channel, 400 km wide at its narrowest point. Madagascar, an island about 1,600 km long, excludes swell from the Indian Ocean from much of the Mozambique coast, except in the extreme south (below latitude 25° S) and in the extreme north, towards Cape Delgado. The southern two-thirds of the Mozambique

coast receives waves generated by southeast trade winds as well as ocean swell from this direction, and these arrive parallel to the coastline between Beira and Angoche and  between Maputo (formerly Lourenço Marques) and Inhambane, where they shape sandy beaches. South of Maputo and on the east-facing coast between Inhambane and Beira, these southeasterly waves arrive at an angle to the coastline, and produce northward longshore drifting. The central part of the coast has prevailing onshore winds that produce short, steep waves across the Mozambique Channel, which are erosive, particularly at high tide, whereas at low tide they form spilling breakers, which build up the beaches. Southerly gales also produce strong wave action, but in the Beira region there are alternations  of northward longshore drifting produced by SE waves and southward longshore drifting, produced by NE waves. Cyclones in the Mozambique Channel can generate waves up to 10 m high, which persist, gradually diminishing, for several days. From Angoche north to the Tanzanian border at the Rovuma River, the coast is dominated by waves generated by the monsoons, with seasonal variations in longshore drifting, northward in winter and southward summer, and no net longshore movement. The South Equatorial Current divides off Cape Delga­ ­do into north and south streams, the south stream becoming the warm Mozambique Current, which passes close to the central coast, attaining a velocity of more than 6 km/h in the October–February, when the northeast monsoon is blowing. Further south, there are large counter-currents along Sofala Bay and near Maputo. The Mozambique Current passes through the Mozam­ bique Channel and converges with the South Equatorial current that flows down the east coast of Madagascar to form the Agulhas Current off South Africa. Tide ranges are generally large because of the broad shallow continental shelf and the Mozambique Channel. Maputo has 3.6 m, but Inhambane only 1.0 m, Beira has 5.6 m, Chinde 3.6 m, and Pemba 3.5 m. The intertidal zone is wide, particularly on the coast between the Save River and the Buzi River at Beira, where mangroves back extensive areas of sand and mud exposed at low tide. Mangroves are extensive on the Mozambique coast, particularly in

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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sheltered bays and on the shores of the Zambezi delta, but their area has been much reduced by cutting for firewood, building, charcoal, and clearance to for salt-making la­­ goons and shrimp ponds. Coral reefs occur off much of the coastline, but fringing reefs exist only north of Porto Mocambo.

2.  The Coastline From the South African border, the southern coast of Mo­­ zambique consists of sandy beaches, some with beach rock platforms (>Fig. 15.1.1), backed by dunes (>Fig. 15.1.2) and Quaternary dune calcarenites that run out from de southern border, in Ponta do Ouro, to Cape Santa Maria on the Machangulo Peninsula and continue to the Save river mouth, including both the mainland and island coasts. Outlying Inhaca Island is the southern limit of coral reefs, although small coral communities extend further south, into South African waters. Fringing reefs are backed by sandy beaches, wide at the NE end. Behind is Delagoa (Maputo) Bay, and the port and capital city of Maputo occupies a cliff-edged promontory on the northern side of  the estuary of the Matola, Umbeluzi, Tembe, and Ma­­puto Rivers, which provides a natural harbour. It consists of  a multiple semi-closed estuary with complex sediment dynamics (Achimo 2000). From Lake St Lucia in South Africa north across the border into southern Mozambique is a coastal dune fringe

up to 150 m high that ends at Cape Inhaca (>Fig. 15.1.3). Along the coast north of Kosimeer Lagoon to Inhaca is a platform of dune and beach calcarenites, and there are some cliff sections. Baia de Maputo (26° S) is much like Hervey Bay in Queensland, Australia, with shoals and tidal channels. The Machangulo Peninsula rises to 86 m, and Inhaca Island to 76 m and a peak of 114 m. Inland is a broad lowland with old S-N dune ridges and swales threaded by the Rio Maputo and backed by the  steep Lebombo Mountains. The Machangulo Pen­ insula has several lakes. Some are of fluvial origin (meander lakes), others are ancient lagoons perched at different altitudes. To the north, the curving coast has beaches backed by dune ridges and elongated lagoons, some of the ridges quite high (up to 190 m), on to the Inhambane Peninsula, which stands about 60 m above sea level. As the coast swings northeast from Maputo, it becomes exposed to ocean swell from the southeast, and there is a long sandy surf beach backed by dunes at Macaneta Point, where a spit has grown southward to deflect the mouth of the River Nkomati (or Incomati). To the northeast, the beach becomes a dune-capped barrier in front of Uembje Lake, a semi-closed coastal lagoon at Bilene, that nowadays is maintained artificially opened and is interrupted by the mouth of the Limpopo River, where are extensive swamps but no protruding delta. The beach resumes past Xai-xai and Chongoene, and becomes a barrier in front of a series of lagoons: Lake Inhampavala, Lake Quissico, and Lake Poelela. At Zavora beach, where there is a narrow

⊡⊡ Fig. 15.1.1 The sandy beach of Xai-Xai, sheltered by a beach rock platform, 1999.

Mozambique

15.1

⊡⊡ Fig. 15.1.2 Cliffed dunes along the southern coast of Mozambique. Though generally forested, these dunes are bare and eroding on the backshore slopes. (Courtesy T.P. Dutton.)

⊡⊡ Fig. 15.1.3 Cape Inhaca, north end of Inhaca Island, cut into dune calcarenite. In the foreground, there is a beach rock platform with bio-karstic microforms, and is covered by white sand, 2004.

fringing coral reef, the coast turns NE past another lagoon, then northward from Jangamo to the dunes of Barra Falsa Point. Barra Falsa Point marks the end of a peninsula that shelters Inhambane Bay. To the north is Linga Linga Point,  site of an old whaling station, and an east-facing coast extending to Pomene and Barra Falsa Point, beside a ­mangrove-fringed estuary, and on to Cape S. Sebastiao (Tinley 1985). Several elongated promontories cut into dune calcarenite, separate low areas with lakes and swamps, and their trend continues in the Bazaruto Archipelago to the north, a Na­­tional Park comprising five

small islands (Bazaruto, Ben­­guerua, Santa Carolina, Magarupe, and Bangue) with three systems of dunes connected to inter-dune lagoons  (>Fig.  15.1.4) and to reeffringed sandy beaches and longshore spits. The mainland coast northward is interrupted by small branched estuaries with swampy shores. A narrow spit runs out from Inhassoro to Bartolomeu Dias, between the Govuro river estuarine mangroves and the sea. It has been breached in places by storm waves (Moreira 2005). The sandy sediment migrate to the north, across the conti­ nental shelf, and reach the delta of the Save river. More branched mangrove-fringed creeks and small islands back

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⊡⊡ Fig. 15.1.4 Transgressive dunes and lagoons in central Bazaruto island. Sandy sedimentation along the eastern lagoonal coast, 2004.

Sofala Bay, where there are extensive shoals of fluvial sand. There is a large gap in coral reefs in this bay, because of the turbid nearshore waters. Northward past the old port of Sofala is the Buzi River, then Beira at the mouth of Pungoe River. These and the Save River are subject to major flooding, and deliver large loads of sand and silt to the coast and  continental shelf (Tinley 1985). In February–March 2000 and 2003, there were major floods in southern Mozambique as the result of intense tropical cyclones Eline and Japhet, respectively, that caused severe erosion of beaches. Cyclone Japhet generated waves 8–10 m high that caused a retreat of 11 m in three days on the eastfacing beaches of Bazaruto and Benguerua islands and the mainland coast of Vilankulos, and very strong winds that caused the destruction of mangroves and dune woodland, driving the dune crests landward (Moreira 2005, 2006). From Beira, a low-lying swampy and mangrove (a total of nine species of trees) coast runs NE to the large Zambezi delta. The growth of this delta has been marked by the formation of parallel beach and dune ridges separated by swampy swales (>Fig. 15.1.5), and progradation is continuing on the southern shore. The deltaic coastline extends to Quelimane, which stands beside the mouth of Qua Qua River, a former distributary of the Zambezi, which became silted in the nineteenth century (Tinley 1985). Beyond this are segments of sandy beach at Madal and Zalala, separated by swamp-edged tidal channels. Black sand beaches (containing rutile and ilmenite) occur along the coast north of the Zambezi River mouth. Erosion

⊡⊡ Fig. 15.1.5 View across the Zambezi delta showing parallel dune ridges and swales occupied by fresh water marshes, with some patches of Barringtonia swamp forest in the wetter areas. (Courtesy K.L. Tinley.)

Mozambique

has been prevalent here, indicated by fringes of undercut and dead mangroves, slumped dune vegetation, depleted beaches, and some breaching of vegetated barriers separating swamps from the sea. The continental shelf narrows past the fishing port and colonial tourist resort at Pebane, but toward its outer edge is a chain of islands including Fogo Island, Casuarina Island, and Mafamede Island, where the beaches are nesting sites for green turtles. Known as the Primeira and Segunda Islands, these are isolated kidney-shaped coral reefs with the indentations on the lee (western) side. Coral reef platforms are exposed at low tide, and are surmounted by islands (cays) of coralline sand and reef gravel, deposited by waves refracted over and round the coral reefs. The islands stand at varying heights above high tide level. Some are bare sand and gravel, others have grasses and shrubs and some are densely vegetated. Angoche is an old port on one of several estuarine inlets along the coast to Point Bajone. Tertiary basalts outcrop near Angoche and to the north in Memba Bay and Porto Mocambo. This is a larger bay, and on its northern side is Mozambique Island, linked by a bridge to the mainland. This island is a UNESCO World Heritage Site because of its many historic buildings, including the large Fort of S. Sebastiao at the northern end and the small fortress of S. Lourenço, built on a small islet of beach rock. Both fortresses show high erosion rates, as seen also on the eastern and NE coast of the islands (Moreira 2005). Fringing coral reefs are extensive along the coast north from Porto Mocambo, and beach rock exposed on some depleted beaches showing weathering forms including bioerosion (Moreira 1996). There is a tombolo at Cabaceira Pequena, where the coast turns north along a series of ⊡⊡ Fig. 15.1.6 The emerged reef platform on Goa Island is 4 m above sea level, 2003.

15.1

headlands of Cretaceous and Tertiary rock between bays cut out along fault lines, mainly SW-NE, connected with the graben of Mozambique Chanel (Lächelt 2004). These have also influenced the orientation of submarine canyons in the continental shelf, which is here extremely narrow, the sea floor plunging to over 2,500 m deep within 30 km of the coast. Tertiary basalts outcrop in Nacala Bay and Memba Bay, and in Lurio Bay there are prograding beach ridges south of the Lurio River mouth. Pemba stands on a peninsula beside a large bay. To the north, the coast is formed by the sedimentary formations of the Rovuma Basin, mainly Cretaceous sandstones and Quaternary coralline conglomerates. The deeply indented coast has emerged coral reef headlands forming rugged cliffs 3–8 m high, with extensive aprons of fringing coral. The coast is bordered by the Quirimbas Archipelago, a chain of coralline islands that emerge from the reef platform of Cabo Delgado Province, and are partly covered by sandy and silty sediment. The inner side of the reef is shallow, with dense mangroves. The outer side of the reef platform is narrow and limited by a steep submerged slope  bordering the Mozambique Channel. The islands are deeply bio-karstified (Moreira 1996), with an intertidal net of channels, along which mangroves grow. Emerged platforms surround the islands at different altitudes (2, 4, and 6 m) as evidence of ancient sea level stands (>Fig. 15.1.6). These platforms contain bio-karstic pools, filled with red clay and fine sand of aeolian origin that form small dune. They are similar to those on the mainland coralline coast of Cabo Delagdo, where the karstified platforms are backed by sharp undercut cliffs (>Fig. 15.1.7). Some islands are ­connected to the capes of the mainland coast by sandy tombolos that locally cover beach rock or reef formations.

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⊡⊡ Fig. 15.1.7 The edge of the platform on Pemba (4 to 5m), bio-eroded with lapies, and a deep notch cut into hard coral conglomerate. Wimbe beach, 2004.

The northern limit of the Mozambique coast is the delta of the Rovuma River, which forms the border with Tanzania. It has a retreating coastline because the sandy sediment budget is negative.

References Achimo M (2000) Sediment types and dynamics of Maputo Bay and Maputo Estuary. Mozambique, Stockholm University, p 57 Lächelt S (2004) Geology and mineral resources of Mozambique. Maputo, Direcção Nacional de Geologia, p 514

Moreira MESA (1996) Bio-erosional forms in beach rock reefs on the coasts of Mozambique and Brazil. Z Geomorphol 106:151–168 Moreira MESA (2005) A Dinâmica dos Sistemas Litorais do Sul de Moçambique durante os Ultimos 30 Anos. Finisterra, Rev Port Geogr 79:121–135 Moreira MESA (2006) Tropical cyclones and their effects on Mozambique coastal dynamics. In: Kelletat D (ed) Guide Book of Coastal Commis­ sion Project Field Symposium on IGCP Project, p 495 Tinley KL (1985) Mocambique. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 669–677

15.2 Tanzania

1.  Introduction The coastline of Tanzania extends from about 4°30' S–10°30' S, a distance of just over 725 km (Alexander 1985). This general outline of the coast appears to have been established by faulting during the Palaeozoic, but local movements may have continued into the late Pleistocene and Holocene. This faulting has occurred within a narrow belt of Neogene sedimentary rock extending along the entire coast of Tanzania (>Figs. 15.2.1 and >15.2.2). The sedimentary rock basin is generally separated from the crystalline basement rock of the interior by faulting, and lie in the down-faulted basin. Evidence from deep borings at a number of places indicates that there has been considerable subsidence of the coast during the Tertiary. Although faulting has blocked out the basic pattern of the Tanzanian coast, the geomorphological details have been determined by fluvial sediment supply, longshore currents generated by  waves and tides, and the growth of mangroves and coral reefs. Marine terraces occur at many places along the Tanzanian coast, indicating a late Pleistocene uplift superimposed on the general crustal subsidence. Two terraces with local coral veneers, about 20–25 and 40–50 m high, occur at Dar es Salaam, Lindi, and the Kilwa area. Coastal and fluvial geomorphology evidence suggests that there are four former marine levels in the vicinity of Tanga, at 2–3, 4.5–6, 24–27, and 41 m, the latter three being most prominent. All are believed to represent sea level changes, with the lower two most likely being related to minor Holocene sea level fluctuations. A single coral-covered terrace, 8–10 m in elevation, extends north from Tanga to the Kenya border (Alexander 1968). In places, the terrace platforms have been arched by crustal movement. There is evidence for tectonic movement along the coast south of Lindi, which has tilted the two marine terrace platforms in that area. Three marine platforms occur between Pangani and Dar es Salaam (Alexander 1969). Morphological evidence

and radiocarbon dating suggest that the middle terrace may have been occupied by two late Pleistocene high stands of sea level. Many of the river valleys along the coast were incised 40  m or more below present sea level in Late Pleistocene times, and their mouths have been submerged by the sea level rise that culminated in Holocene times to form estuaries. The climate of the Tanzanian coast is monsoonal, with NNE winds from November–March and SE winds from April–October. There is a short rainy season in November, and a longer one between March and May, but in the south there is a pronounced dry season followed by wet weather from November to May. During the rainy season (especially March–May), Tanzanian rivers transport large amounts of sediment to the coast. The rivers flow into two distinct environments. The Rovuma, Lukuledi, and Pangani flow through narrow estuaries and into deep water where most of their sediment is lost. The Rufiji and Wami Rivers discharge into the shallow water in the lee of Mafia Island and Zanzibar, where they have formed deltas. Sediment from the Rufiji, Ruvu, and Wami Rivers is carried northward by longshore drift and nearshore currents, which prevail throughout the year although the rate of flow changes seasonally. Mean spring tide ranges on the Tanzanian coast are just over 3 m (3.3 m at Mtwara and Dar es Salaam) and 3.7 m at Zanzibar Town. Mangrove forests grow in sheltered bays and estuaries, and are abundant along the shores of the Rufiji, Ruvu, and Wami deltas. They also occur in some more exposed locations along the coast, particularly along the open coast between the Rufiji and Rovuma rivers and north of the Pangani River where they are frequently found along the open shore, although in places their growth is impeded by wave action (Alexander 1966). Coral reefs (fringing and offshore) and associated depositional islands are found along much of the Tanzanian coast and around the islands of Mafia, Zanzibar, and Pemba. Beaches are extensive, and consist largely of coralline sediment derived from the reefs, which generally front

Edited version of a chapter by C.S. Alexander in The World’s Coastline (1985: 691–695). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 15.2.1 The Tanzanian coast north from the Rufiji River delta.

Tanzania  

⊡⊡ Fig. 15.2.2 The Tanzanian coast south from and including the Rufiji River.

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⊡⊡ Fig. 15.2.3 Beach erosion near the Africana Hotel, Kunduchi. (Courtesy Geostudies.)

them, but in the vicinity of river mouths (particularly on delta shores) quartzose beach sands are of fluvial origin. At low tide, a gently sloping foreshore is exposed, with sand bars and muddy areas, often passing into a fringing coral reef (Alexander 1966). Beach ridges occur in several sectors, particularly on  the northern side of the Rufiji delta, between the Pangani and Ruvi Rivers and in northeast Tanzania (Alexander 1969). In many places, the old beach ridges are now being truncated by marine erosion (Alexander 1966) (>Fig. 15.2.3). This erosion may have been going on for some time, for erosion was reported at widely separated locations at the beginning of the twentieth century. Beach erosion appears most active, at least between Dar es Salaam and Bagamoyo, at the onset of the monsoons in May and November. North of the Moçambique border, which follows the Rovuma River, the coastline is deltaic with a sandy beach fringe, past the new port of Msimbati. A fringing coral reef begins, interrupted at the head of Mtwara Bay where the old coastal town on Mtwara stands, and again at Sudi Bay, the elongated estuary of Mambi River. Between the river mouths are segments of emerged coral reef. Lindi stands on the sandy north shore of the Lukuledi River estuary and to the north the emerged coral reef is cliffed and bordered by a modern fringing reef, interrupted at the mouth of the Ushingi River in Mchinga Bay and again in Kisware Bay. To the north, the emerged reef breaks into islands sheltering Sangarungu Haven, including Kilwa Kisiwani island and Songa Mnara island.

On the north side is the old town of Kilwa Masoro, and the Matandu River flows into the lagoon-like bay. North of Kilwa Kisiwani, the continental shelf widens and a series of coral reefs curves out along its edge toward Mafia Island. Several have cays, as at Songo Songo. Emerged coral continues along the coast northward, the cliff declining into a sandy beach, which runs on round the large Rufiji River delta, built in the shelter of Mafia Island. The river breaks into distributaries, which open to the sea through swampy inlets. North of the Rufiji delta, the coast is low-lying and alluvial, but fringing reefs border Koma Island and Kwale Island offshore. At Ras Pembamnasi, fringing reefs reappear on the mainland coast, again fronting an emerged coral reef that has cliffed sectors. These continue past Ras Kimbiji to the inlet of Dar es Salaam Bay, fed by the Msinga River. Offshore the Makatumbe Islands are flattopped, cliff-edged slabs of emerged coral (>Fig. 15.2.4). Dar es Salaam, the capital of Tanzania, occupies a broad promontory of emerged coral on the northern side, with reddish cliffs and beaches. To the north, the coast passes into the shelter of Zanzibar Island, and there is relatively low wave energy  along the shores of Zanzibar Channel. Beyond the  cliffed peninsula of emerged coral on the Ras Kankadya  is Msasana Bay, which has sandy beaches backed by numerous parallel Holocene beach ridges and  fronted by a  wide intertidal zone where exten­­ sive  orthogonal sand bars are exposed at low tide (>Fig. 15.2.5).

Tanzania  

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⊡⊡ Fig. 15.2.4 Makatumbe, an emerged coral reef island off Dar es Salaam. (Courtesy Geostudies.)

⊡⊡ Fig. 15.2.5 Sand bars in the intertidal zone near Kunduchi. (Courtesy Geostudies.)

Msasana bay is sheltered by Bongoyo, Pangavini, and Mbudya Islands, emerged flat-topped coral reefs with cliffed margins showing notches cut largely by solution, but with some local abrasion ramps. These are bordered by low-tide shore platforms cut in the emerged coral, with algal ridges and some patches of coarse coralline gravel, and corals growing at the seaward margin indicate that a fringing reef could be built up to the same

level as the shore platform here (>Fig. 15.2.6). On the inner (western) shore of Mbudya Island, the emerged coral has been dissected into irregular karstic topography (>Fig.  15.2.7). There is a cuspate sandy foreland with successively formed beach ridges indicating its ­pattern of growth, but lines of beach rock show that it was formerly larger, and has been cut back on the northern flank.

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⊡⊡ Fig. 15.2.6 Cliff and notch in emerged coral limestone on Mbudya Island, showing low-tide shore platform. (Courtesy Geostudies.)

⊡⊡ Fig. 15.2.7 Karstic shore topography on emerged coral limestone on the west coast of Mbudya Island. (Courtesy Geostudies.)

Small streams descend valleys incised into the emerged coral terrace and the parallel beach ridges, and open to mangrove-fringed inlets interrupting the sandy beach. At Kunduchi, the beach ridges form a barrier spit on the seaward side of Perembji Creek, which has been deflected northward by longshore drifting. Off the creek mouth is an intertidal fan of outwashed coarse rippled sand. Although this is a low wave energy environment, the

­ reviously prograded beach ridge fringe is being cut back p by marine erosion, possibly as the result of recent coastal subsidence (>Fig. 15.2.8). Cliffs fronted by beaches continue northwest from Bahari, past the headland at Ras Kiromani and along Unonio Bay. Longshore drifting has built a spit at Mlingatini, sheltering a bay into which flows the River Luale. To the northwest is the coastal town of Bagamoyo,

Tanzania  

   15.2

⊡⊡ Fig. 15.2.8 Beach ridge crests on the coast north of the Rufiji River. The pattern suggests at least two phases of deposition, separated by periods of erosion. Erosion is now truncating the youngest set of ridges.

and the sandy beach ridges continue past the Ruvu River and round the swampy Wami delta, and along the Sadani Game Reserve. The emerged coral forms a narrow coastal terrace with low cliffs along much of the northern coast of Tanzania. Wave energy increases as the coast passes beyond the shelter of Zanzibar Island and the continental shelf narrows northward. A chain of coral reefs resumes off Mkwaja, and continues past Karongo Island and Yambe Island. Pangani River reaches the sea by way of an estuarine inlet, and the coast then passes into the lee of Pemba Island, with wave energy diminishing over the Pemba Channel. To the north, the emerged coral coast becomes indented. Tanga is a port beside a bay with swampy shores, and low cliffs run behind Manza Bay and around Kirui Island, a detached segment of emerged coral, to the Kenya border. Offshore the chain of coral reefs continues near the edge of the continental shelf.

2.  Mafia, Zanzibar, and Pemba Islands The islands of Mafia, Zanzibar, and Pemba show evidence of several terraces in emerged Pleistocene coral limestone and Neogene sedimentary rocks, along with indica­­ tions of a relatively recent subsidence forming inlets and valley-mouth bays. The Neogene sediment, ranging from Miocene to Cretaceous in age, are calcareous sands and clays of marine origin, and more than 3,300 m thick. The three islands are of similar size, but while Mafia and Zanzibar lie on the continental shelf, Pemba is separated from the mainland by a 700 m deep channel that is probably the result of graben faulting. The eastern coasts, where deep water comes close inshore, are thought to be bounded by faults. Mafia Island has a low plateau (30–45 m) on Neogene rocks on the western side, separated by flat-bottomed, swampy valleys that may indicate recent subsidence of

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the  island. Historical subsidence is indicated by remains of  medieval structures on the west coast that now lie between high and low tide levels. Along the east coast is a belt of emerged coral reef about 1.5 km wide that stands about 3–5 m above sea level. The emerged reef is covered by poorly bedded aeolianite of variable thickness to a maximum of 21 m. Inland, a smooth area 15–18 m in elevation may represent a second marine terrace. A low sea cliff extends all along the eastern side of the island; to seaward is a narrow, modern fringing reef that extends around much of the island, and widens considerably on the south coast. To the north, the emerged reef runs out to a sharp point, and to the south east it is breached by broad Chale  Bay, then continues in Juani and Jibondo Islands. Mangroves grow in bays and inlets on the northwest coast, but the east coast is cliffed, with fringing reefs in and south of Chale Bay. Zanzibar is dominated by emerged reef terraces, except for two linear corridors of Neogene rock in the northwest. The reef terraces indicate Pleistocene uplift of up to 50 m. There are a few small rivers, draining mainly to inlets on the northwest coast. On the east coast, exposed to Indian Ocean swell, the emerged coral terraces are cliffed, and bordered by wide modern fringing reefs, extending round the north cape, Ras Nungwi. The east coast cliffs are interrupted by Chwaka Bay, which has mangroves on its southern shore and along the western side of the bordering Ras Michamvi Peninsula. To the south, the cliffs are fronted by sandy beaches, as at Jembiani, and Makunduchi lighthouse stands on a coral headland. The port of Zanzibar Town stands on the west coast. To the north are several coral reefs with cays, as at Chumbe Island, and Tumbatu Island is reef-fringed. There are sandy beaches behind fringing reefs, as at Mangapwani in the northwest and Kizimkazi in the southwest.

Pemba Island is dominated by Neogene formations, with an emerged coral terrace along the east coast. The east coast is relatively straight, while the west coast has intricate bays and inlets. Valleys descending to these have flat, swampy alluvial floors, aggradation of which may result from a slight recent subsidence of the land. Most of the east coast of Pemba Island consists of an  emerged coral reef about 5–6 m high. Also, there is scattered evidence of a second reef platform at 9 m above sea level. The emerged coral reef is cliffed, and bordered seaward is a fringing reef 200–500 m wide. It extends from the northeast cape, Ras Kiuyu, southward past Kojanil, a bay sheltered by an emerged coral island. Mangroves occur in inlets and on the western shores of islands. Chake Chake Bay is a long inlet opening to the north coast of Pemba Island. Emerged coral terraces also form the northwest peninsula out to Ras Kigomasha, and elongated reef islets Uvinje, Fundo, and Njaq to the south. Mangroves fringe the shores of bays and inlets on the west coast, as in Jombangome Creek, and Wete is a town on the north shore of a western bay. To the south is the port of Mkoani, and a chain of emerged coral islands extends to Kisuwani.

References Alexander CS (1966) A method of descriptive shore classification and mapping as applied to the northeast coast of Tanganyika. Ann Assoc Am Geogr 57:128–140 Alexander CS (1968) The marine terraces of the northeast coast of Tanganyika. Z Geomorphol Suppl(Bd.) 7:133–154 Alexander CS (1969) Beach ridges in northeastern Tanzania. Geogr Rev 59:104–122 Alexander CS (1985) Tanzania. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 691–701

15.3  Kenya

Francis Ojany

1.  Introduction The coastline of Kenya is about 530 km long. Essentially, the whole length of the coast forms part of the Lowland Plain (Ojany 1966, 1973), which has resulted from downwarping of the coastal basin (in Palaeozoic times), marine transgression (in Mesozoic times), and a fluctuating sea level (during Quaternary times). A number of small faults have been mapped along the coastal belt. Of these it is probably the inferred continuation of the Ruvu-Mombasa faul that has had the most significant effect on the relief of the area. Spring tides at Kilindini have a maximum range of 4.0  m, with an average for most months between 2.5 m and 3.6 m. Malindi has a tide range of 2.9 m.

2.  The Kenyan Coastline The southern half of the coast is quite different from the northern half. Up to Malindi the shore sediment are mainly marine and lagoonal in origin (Abuodha 2004). In the center the Ras Ngomeni coral and coquinoid island has been joined by a sandy isthmus to the mainland, thus forming a tombolo. In the north the shore sediment have come mainly from continental and deltaic (or estuarine) environments and are dominated by Quaternary sediment, exemplified by such large sand dunes as Mlima Kitanga Tangani (next to Ungwana Bay) and those to the north up to Ras Biongwe. The Mlima Kitanga Tangani dune sands have formed high outer (rising to 52 m) and inner barriers. These barriers have caused a deferred junction at the mouth of the Tana (here called the Ozi) and the other streams in the area. The dunes are responsible for the constant shifting (or migration) of the mouths of these rivers. In the Lamu archipelago, both the Mongoni and Dodori rivers bring considerable sand and other alluvial material from upstream; they have supplied the rather complex pattern of islands, inlets, channels, and bays in the area. At Malindi there has been progradation south of the Galana River, forming a very wide beach in front of a former coastal resort (>Fig. 15.3.1).

Coral reefs on the Kenya Coast are typical of an equatorial environment. Live coral polyps occur on the present fringing reef. Coral development has been attributed to the growth of polyps up from a marine platform at least 60 m below present datum during the Kamasian (Mindel Glaciation) Age. Coral growth is more vigorous in the south, as shown by the more or less continuous fringing reef. Hori (1970) explained this occurrence in terms of thicker coral development in the Mombasa area but Ase (1981) emphasised the predominance of emergence in the north as opposed to submergence in the Mombasa area. Terrigenous sedimentation, more importantly in the north, tends to impede coral growth. The fringing reefs of the Mombasa area pass northward into nearshore reefs backed by a shallow boat channel (Bird and Guilcher 1982). They show algal networks (>Fig.  15.3.2) and are backed by cliffs cut into emerged Pleistocene coral limestone (>Fig. 15.3.3). The effects of a changing sea level during the Quater­ nary (Pleistocene and Holocene) have left significant features along the Kenya coast. Emerged platforms, beaches, caves, stacks, and nickpoints are all well preserved along the costline. The Changamwe Terrace is at about 45–70 m, the Upper Mombasa Terrace at about 15–37 m, the Lower Mombasa Terrace at about 10 m and the Shelly Beach Terrace at about 5 m. All these features can be seen in the Shelly Beach area, just south of Mombasa, with some excellent examples of abandoned cliffs and former islands. There is also a highly developed 36 m marine platform in the Lower Kwale area which is extensively used to grow sugar cane. The full sequence of events that brought about these changes started back in Pliocene times when the 107 m nickpoint was formed. The final episode did not come about until the present aggradation and silting had occurred. These most recent events include the drowning of the rivers that flowed through the present Port Reitz and Tudor to form Mombasa Island and the other creeks along the coast. The present mangrove swamps and trees then became established as the sea level rose to the present datum. Hori (1970) gave some radiocarbon dates for rock

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 15.3.1 Prograded sandy beach at Malindi. (Courtesy Geostudies.)

⊡⊡ Fig. 15.3.2 Fringing reef with algal ridge networks at Nyali. (Courtesy Geostudies.)

samples from the Nyali, Shelly Beach and Tiwi areas. The ages of the samples ranged from 26,500 years bp to 2,820 years bp. The incision of river valleys to form the present Mombasa and Kilindini harbours created excellent deep navigable channels. The relationships between shore displacement, wave action, fetch and coastal aspect are important for

­ nderstanding the relative erosion rates and the general u orientation of the coastline. The seasonal reversal of the monsoon winds has also played a role. In the extreme south, the SE waves arrive at right angles to the coastline,  while to the north they come in at an angle and ­generate northward longshore drifting, most obvious in Ungwana Bay.

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⊡⊡ Fig. 15.3.3 Cliff and notch in emerged coral limestone, Nyali. (Courtesy Geostudies.)

References Abuodha J (2004) Geomorphological evolution of the southern coastal zone of Kenya. J Afr Earth Sci 39:517–525 Ase LE (1981) Studies of shores and shoreline displacement on the southern coast of Kenya. Geogr Ann 63:303–310 Bird ECF, Guilcher A (1982) Preliminary observations on the modern fringing reefs of Kenya and the associated shore forms (in French). Rev Geomorphol Dynam 31:113–125

Hori N (1970) Raised coral reefs along the southeastern coast of Kenya, East Africa. Geogr Rep Tokyo Metrop Univ 5:25–47 Ojany FF (1966) The physique of Kenya a contribution in landscape analysis. Ann Assoc Am Geogr 56:183–196 Ojany FF (1973) Kenya: a study in physical and human geography. Longman, London

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15.4  Somalia and Djibouti

1.  Introduction The Somali coast is about 3,200 km long, of which 2,100 km face southeast to the Indian Ocean and 1,100 km face north to the Gulf of Aden. The coast has a few natural harbours, is fronted for much of its length by coral reefs and is backed throughout by desert or dunes. Coastal access is  thus difficult, and understanding of coastal features remains limited. The coast may be divided into three zones. First, from Raas Jumbo on the Kenya border (1°39’ S, 41°36’ E) to near Gifle (7°28’ N, 49°40’ E), the 1,500 km generally low coast is backed by massive Quaternary dunes and fronted by coral reefs. Second, from Gifle to Ras Asir (11°49’ N, 51°15’ E), the 600 km coast is 100–300 m high and often steeply cliffed, for example on Ras Hafun (10°27’ N, 51°24’  E), Africa’s easternmost point. Third, from Ras Asir to the Djibouti border west of Seylac (11°15’ N, 43°30’ E), the guban (sunburnt plain) coast along the Gulf of Aden is partly cliffed, and partly backed by alluvial plains, sebkhas and marine terraces. The outlines of the Somali coast were defined by Cainozoic tectonic activity, while the details were formed by Quaternary erosion and accretion. During Quaternary times there was variable uplift along the north coast of Somalia, as indicated by marine terraces. Inland, the north-facing escarpment of the Golis Ranges has been much dissected along fracture zones, exposing Pre-Cambrian basement rocks and deformed Mesozoic sediment. Farther south, Somalia is essentially a plateau surface that slopes south and east from the northern mountains and the Ethiopian Rift toward the Indian Ocean. The east coast of Somalia was apparently outlined by Tertiary faults. The Somali coast is dominated by a reversing monsoon system, which finds expression in surface winds, wave action, ocean currents and onshore climates. Four seasons are commonly recognised. From late December to March, the jilaal season is dominated by the NE monsoon, and hot dry dusty winds blow along the north and east

coasts at velocities of 15–30 km/h. From March to May, the transitional gu season, the winds change direction and bring rains from the ocean. The long hagaa season, beginning in June and continuing through September, is dominated by the SW monsoon with mean wind speeds often exceeding 40 km/h. Along the north coast this is the season of a strong, hot, desiccating offshore wind, the socalled kharif, which reaches maximum velocity around dawn. The air is filled with blowing sand and dust, and temperatures along the Gulf of Aden may reach 50°C. In the south, the cooling winds off the ocean and occasional rains render conditions more pleasant, but northward sand transport is common in the dunes as winds frequently exceed threshold velocities. The dayr season in October and November may again bring rains during the lull between the SW and NE monsoons. Although rainy seasons may be recognised in principle, the actual precipitation is negligible, except in the northern mountains and the extreme southwest. Rainfall on the north coast is only 50–150 mm/year, and the region is persistently hot and humid. The east coast south of Ras Assuad receives as much as 300–500 mm/year and humidity generally exceeds 70%, but with temperatures ranging from 20°C to 35°C semiarid conditions prevail. The coast farther north lies between these two climatic extremes but may sometimes be afflicted by flash floods triggered by downpours in the mountains. Generally, however, few streams reach the sea even under ephemeral flow conditions, and Somalia has only two perennial rivers, the Shebele and Juba in the south, which drain from the wetter Ethiopian Highlands. Infrequent tropical cyclones may penetrate the Gulf of Aden from the Arabian Sea, bringing gale-force winds and torrential rain, generally during the changeover seasons between the dominant monsoons. Land and sea breezes are also common when the monsoon flows are weak. Sea breezes add an onshore component to both the NE and SW monsoons during the daytime, which may explain the orientation of the coastal dunes subparallel with the coast.

Edited version of a chapter by Anthony R. Orme in The World’s Coastline (1985: 697–711). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Indian Ocean coast of Somalia is exposed to swell generated in the Southern Ocean. The continental shelf is narrow, and there are few islands. Swells and wind-driven waves from the south predominate, and longshore drifting is mainly to the northeast. Average wave heights exceed 1.0 m during the southwest monsoon, about 0.5 m at other times. Between Ras Assuad and Ras Hafun, wave heights over 1.5 m occur 20% of the time. Throughout the coast, tides are mixed and spring tides range from 4 to 6 m. Under present climatic conditions sediment discharge from wadis is minimal, except from the Juba River during occasional floods. In ecological terms, persistently high air temperatures above 20°C favour mangrove vegetation in suitable habitats, while sea surface temperatures over 20°C and low turbidity favour extensive coral growth. Tsunamis are rare, but a major one occurred on Boxing Day 2004 as a result of an earthquake south of Sumatra. After an initial withdrawal, a series of waves up to 3 m high broke on the coast of Somalia, causing extensive sea flooding and structural damage. On several sectors sand was washed onshore by these waves, and deposited on and behind pre-existing beaches.

2.  The Coast of Somalia For 200 km from Raas Jumbo in the extreme south to the Juba River mouth, the coast is less harsh and more varied than elsewhere in Somalia. Dunes of various ages are covered by modified scrub and open woodland, but their continuity is broken by the valleys of seasonal streams, some of which may reach the sea in long inlets, for example the 20 km long Buur Gaabo. Mangroves are common around such inlets and around the Juba River mouth. Avicennia marina is the dominant species, but Rhizophora mucronata, Ceriops tagal and Bruguiera gymnorrhiza are also found. The once lush mangroves have been depleted by centuries of woodcutting. Offshore, intermittent coral reefs are found, while the numerous Juba Islands and islets appear to be remnants of raised reef formations. From the Juba River for 1,300 km northeast to Gifle the coast is fringed by extensive coral reefs. No rivers reach this coast, natural harbours are lacking, and the principal ports of Baraawa (Brava), Marka (Merca) and Mogadishu have always been difficult to approach. Occasionally, low headlands are of emerged Quaternary reef limestone (a 6–9 m terrace being locally prominent), dune calcarenites and beach rock. Otherwise the coast is dominated by dunes of various ages, the oldest of which rise to over 200 m above sea level. In places, these dunes rest upon coral reefs. Between

War­shiikh and Baraawa the Shebele River has been deflected behind coastal dunes to flow southwest for as far as 400 km towards the Juba River. During rainy periods, and following exceptionally heavy rains in the Ethiopian Highlands, the Shebele may flow through to the Juba, but most of the time its water disappears into the sand or swampy areas along a low corridor behind the dunes. Newly built dune fronts and backshore sand terraces along the coast are colonised by the sprawling beach vine Ipomoea pes-caprae, with grasses and occasional bushes (Pardi 1976). Above the highest tides there are scrub thickets, with rushes in marshy sites. Mangroves grow locally along this coast but are not well developed. The older dunes along this coast have been destabilised by overgrazing of their vegetation cover by cattle, goats and camels (Orme 1982). Destabilisation has been followed by extensive sand movement and by gullying of the more indurated sands on the moist seaward face of the dunes. Around the major settlements, refugees have swelled the local population, and tillage has been practised on the dunes, leading to severe erosion. The reactivation and erosion of the older dunes over relatively large areas poses an enormous problem and attempts have been made to stabilise this. Much attention has focused on the Marka area, where erosion gullies have been erased by bulldozers and their red sandy sediment sown with vegetation, and where attempts have been made to arrest the migration of the dune front on to the valuable farmlands of the Shebele Valley, where some settlements had been buried beneath drifting sand. From Gifle to Ras Asir much of the 600 km coast is cliffed, cut into horizontal or gently folded Cainozoic carbonate rocks. Ras Illigh is a rocky terraced headland rising 145 m above the sea. From Eyl to Ras Mabber the bold rocky coast has sea cliffs rising 75–120 m above sea level. On the Hafun peninsula a former island is connected to the mainland by a 25 km tombolo (>Fig. 15.4.1). It consists of Oligocene marine limestones, including corals, overlain unconformably by Miocene sandy limestones and compact coral limestones forming cliffs up to 203 m high. At Ras Asir, Cainozoic sedimentary rocks form sheer 300 m high cliffs. Between these cliffed sections, there are some marine terraces, partly blanketed by blown sand, and some inlets with mangroves. The Daror and Nugaal rivers drain large valleys, but under present climatic conditions their flows are localised and erratic, and the valleys have never carried water through to the coast in recent times. As in other parts of the Horn of Africa, the size of these valleys and the extensive legacy of fluvial deposits are attributed to past pluvial conditions. The 1,100 km long guban coast, extending from Ras Asir to beyond Seylac, comprises a coastal plain of

Somalia and Djibouti

⊡⊡ Fig. 15.4.1 The tombolo on the Hafun peninsula. (Courtesy Geostudies.)

15.4

oriented NE–SW, parallel to the dominant winds. Farther east, beyond Raguuda, the coastal plain narrows to less than 10 km and sometimes disappears as basement rocks and later sedimentary formations reach the coast in bold cliffs. Three specialised ecological habitats characterise the coast: mangrove, coral flat and sand dune. The mangrove vegetation is mostly identical in composition to that noted on the east coast, with Avicennia marina being the dominant species. The low raised reef flats are poorly drained and support sparse perennial herbs, and near water sources the dunes support shrubby vegetation.

3.  Djibouti

v­ ariable but generally narrow width, backed by the dissected Golis escarpment, which reaches 2,408 m on Surud, 60 km SSW of Ras Sura. During seasonal rains in these mountains much sediment is discharged on to the lowlands in the form of extensive alluvial fans and bajadas. During earlier Quaternary times along this tectonically active coast, successive episodes of alluviation caused fluvial sediment to become interdigitated with marine terrace assemblages. There is an emerged coral reef about 8 m above sea level, and the surface of a higher terrace at 18 m is littered with corals, echinoids and shells. There are also Pleistocene and Holocene corals in beach deposits 180–280 m above sea level on the Berbera coastal plain (Mason 1962), and similar deposits up to 120 m in the area between Laas Qoray and Boosaaso. Boreholes near Berbera revealed a thick Plio-Pleistocene marine succession, indicating relatively late tectonic uplift in this area. The guban is widest in the west, where it is covered by much fluvial debris and a veneer of dune sands with ridges

The coastline of Djibouti (formerly French Somaliland) extends from the Somali border around the Gulf of Tadjoura to the western shores of the Bab el Mandeb (12°40’ N) at the southern end of the Red Sea. The nearshore shallows extend westward in the Gulf of Tadjoura, past the port of Djibouti, built partly on a spit formation. This part of the coast is backed by rugged hills of red volcanic formations. The Gulf then narrows westward to a strait leading into the almost landlocked Ghoubet Kharab, which occupies a structural depression. North of the Gulf of Tadjoura a hilly coastline with numerous wadi-mouth sand plains declines to a coastal lowland with sandy beaches bordering the Bab el Mandeb, facing the strait of Aden.

References Mason JE (1962) The area North of Hargeisa and Lageruf. Somali Republic Geological Survey Report 7 Orme AR (1982) Field observations along the Somali Coast. Depart­ ment  of Geography, University of California, Los Angeles, pp 1981–1982 Pardi L (1976) Researches on the Somali coast: shore and dune of Sar Uarle. Monit Zool Ital 5:179–193

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15.5 Eritrea

Eric Bird

1.  Introduction The coast of Eritrea is about 900 km long, facing northeast to the Red Sea. It is a hot and arid coast: Massawa has a coldest month (January) with 26°C and a mean annual rainfall of only 130 mm. There are several watercourses, but rivers flow only briefly after rare rain, and many do not reach the sea. Although there are rugged mountains running SE–NW, there are few steep sectors, much of the coast being backed by broad pediments descending from mountain fronts. Tectonic movements have continued in Quaternary times, and there are emerged coral reefs locally along the coast. Winds and waves on the Red Sea are mainly northeasterly and easterly. Tide ranges are small: Massawa has a mean spring tide range of 0.9 m. Salinity is high (sometimes exceeding 40‰) in the Red Sea, evaporation losses being compensated by an inflowing current from the Indian Ocean through the Bad el Mandeb. Coral reefs are abundant, many of them with sand cays. The largest array of coral reefs is in the Dahlak Archipelago seaward of the Massawa Channel. This includes various reefs, such as Entedebir Island, where there has been disruption by faulting so that reef platforms stand at various levels up to 14 m, while the surrounding submerged reefs contain faulted troughs. Mangroves occur locally in sheltered bays and inlets, but are sparse.

This if followed by a long sector where a narrow pediment fringing the Denakil Range is crossed by successive parallel watercourses running out from mountain-front wadis. There are beaches at the mouths of these watercourses, interspersed with low cliffs on reefs and segments of fringing reef. Near Edd the coast is hilly and irregular, and beyond Thio, Amfile Bay has many reefs and cays. The coast then becomes indented, with ridges running out to small promontories including Ras Andadda. The broad Bay of Howakil contains several high islands and numerous reefs, and beyond it the wide Buri Peninsula runs out to the shores of Massawa Channel. The outlying Dahlak Archipelago is a reef complex, and has been designated as a major marine National Park. On the western side of the Buri Peninsula is the deep  Gulf of Zula, a trough bordered landward by high ­mountain ranges. The tectonic depression is bordered by warped and dislocated emerged coral reefs (Lalou et al. 1970). These continue NW behind Massawa, where wave energy is low in the lee of the Dahlak Archi­pelago. To the north, lobate promontories separate bays at the mouths of valleys such as the Wakira, and then the coastline becomes almost straight as a pediment ­widens in front of the high mountains. Several parallel watercourses emerge from wadis and cross the pedi­­ment,  some  ending in small marine inlets, as at Marsa Mubarek, Marsa Deresa and Marsa Berisse, with fringing coral reefs. The pediment coast extends across the border into Sudan.

2.  The Eritrean Coast North of the Djibouti border the Strait of Bab el Mandel is bordered by a peninsula that runs out to Ras Sinhan, sheltering the Bay of Assab, containing Haleb Island and a number of coral reefs and cays. Assab stands in front of a hill from which a ridge runs NW to a point at Ras Dormo.

Reference Lalou C, Nguyen HV, Faure H, Moreira L (1970) Uranium-thorium dating of high levels of coral in the Afar Depression, Ethiopia. Rev Géogr Phys Géol Dynam 12:3–8

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

15.6 Sudan

Eric Bird

1.  Introduction The coast of Sudan is about 650 km long, and runs along the Red Sea. It is hot and arid: Port Sudan has a temperature of 23°C in January and 34°C in July, with a mean annual rainfall of 94 mm. The hinterland is high from south to north, but there is a coastal lowland consisting of broad pediments that descend from mountain fronts. There are several watercourses, but rivers flow only briefly after rare rain, and many do not reach the sea. Tectonic movements have continued through Quaternary times, and there are emerged coral reefs along the coast. Salinity is high (sometimes exceeding 40‰) in the Red Sea, where coral reefs are abundant, many of them with sand cays. Winds and waves on the Red Sea are mainly northeasterly and easterly. Tide ranges are very small: Port Sudan has a mean spring tide range of 0.1 m, which is obscured by sea level fluctuations of 1 m above and below mean sea level, resulting from changes in barometric pressure and onshore and offshore winds. Sparse mangroves grow locally in sheltered bays and inlets.

2.  The Sudanese Coastline From the Eritrean border the coastline is intricate, with headlands such as Ras Kasar formed by segments of emerged coral reef. There are fringing and patch reefs in coastal waters. The wide pediments descending to the coast continue north to Suakin, beyond which there is a large sebkha (a hypersaline flat subject to rare flooding) fed by  watercourses descending from mountain valleys. Off Suakin is Green Reef, one of the few atolls in the Red Sea. Port Sudan is located beside a small inlet known as a marsa or sharm formed by marine submergence of the mouth of a valley. The outlying reefs include another atoll, Sanganeb (Rathjens and Von Wissmann 1933). Several

such inlets are found along the coast to the north, notably at Marsa Darur and Marsa Salak, the pediment ending in low cliffs that include segments of coral reef. In places the coral reef that has emerged is several kilometres wide, very well preserved in this arid environment and backed by the former boat channel, filled with downwashed sediment (Dalongeville and Sanlaville 1981). Ras Abu Shagara is a promontory that is in the form of a large spit extending southward, sheltering Dungunab Bay. To the north, the coast is dominated by coral reefs emerged, across which numerous inlets have been incised, including Marsa Shinab, Marsa Delwin, Marsa Oseif and Marsa Marob. Some of these have narrow  entrances cut through the fringing reef, and become wider landward, where the Pleistocene valley was cut into the softer sediment of the infilled boat channel prior to Holocene marine submergence. At Cape Elba (Ras Hadarba) the coast swings northwestward, the pediment continuing past Halaib. There are numerous coral reefs and cays, as well as emerged reef structures. In the hinterland the mountain front rises steeply to Mount Elba (1,594 m) at the northern end of a high ridge. The coastal plain then widens as the mountains recede behind an enclave of confluent wadis descending to Ras Abu Dara. The coastline is irregular, punctuated by many small marine inlets, including the large branching Marsa Sha’b, where the mouths of Wadi Ibib and Wadi Hodein meet. The irregular coastline continues across the border into Egypt.

References Dalongeville R, Sanlaville P (1981) The marssa of the Sudanese shore of the Red Sea (in French). Bulletin Géographique Languédoc 15:39–48 Rathjens C, Von Wissmann H (1933) Morphological problems in the Red  Sea graben (in German). 183–187 Petermanns Geogr Mitt 79:113–117

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

15.7 Egypt, Red Sea Coast

Anja Scheffers · Tony Browne

1.  Introduction The Red Sea coast of Egypt, including the Gulfs of Suez and ‘Aqaba and the intervening Sinai Peninsula, is about 1,500 km in length.

2.  Egypt Red Sea Coast Ras Hadarba (Cape Elba), just north of the Sudanese border, is a small promontory on a coast facing ENE where a pediment fronts arid mountains. There are intermittent fringing coral reefs, with gaps near the wadis that descend from the mountain front across the coastal pediment. Occasional bays and inlets, here termed sherms, indent the coastline, as at Sherm el Madfa, and mangroves also exist locally. To the north is Foul Bay, with a barrier spit fronting an open lagoon at Berenica and a peninsula running out to Ras Benas, where there are outcrops of Pliocene calcareous grits. To the north, a long section of almost straight coastline parallel to the mountains and their fronting pediments, with wadis and occasional mersas (inlets) and segments of fringing reef, extends past Quseir to Port Sarfaga. Here a larger bay is partly blocked by Sarfaga Island, which has fringing coral reefs and a mangrove park. Port Sarfaga exports phosphates, and the adjacent resort area has several hotels, some with artificial beaches (>Fig. 15.7.1). Soma Bay, to the north, has a wide intertidal zone, which is partly a dead coral reef. The sandy beach at El Gharib has exposures of stratified sandstone (>Fig. 15.7.2), and is backed by dunes about 1 m high, with halophytic shrubs on the northern, windward side (>Fig. 15.7.3). At Ghabour, a very wide pediment descends to the shore, where a beach of shelly sand has emerged, somewhat dissected by gulleys cut by rare pediment run-off. There is a wide beach of algal carpets, fronted by a sandy intertidal platform over planed-off coral reefs, and a few mounds of wind-blown sand.

A prominent arrow spit runs out at right angles to the coastline. Seaward, there are large coral-fringed islands and deeper straits with seagrass vegetation. At Abu Suman, to the north, the shelly beach that has emerged runs out to a spit with an islet bearing dunes and fringed by beach rock. A wadi runs down to Sharm al’ Arab, and beyond the Marsa Abu Mukhaqij Bay is the town of Hurghada. Hotels have been built along this coast, together with piers and swimming pools. The beaches of coarse sand are fronted by an intertidal zone out to nearshore coral gardens, and sand bars move in across the shore. At Hurghada there are exposures of beach rock (>Fig. 15.7.4) and the beach is littered and polluted with oil. North of Hurghada there are mangroves beside a small estuary at El Gouna. Oil rigs and piers and tankers at Ras Shukheir mark the exploitation of a coastal oilfield. Coral reefs become more extensive in the Gerfatin Islands in coastal waters north of Hurghada as the shelf widens out to Shadwin Island and the Strait of Gubal, the entry to the Gulf of Suez. The west coast of the Gulf of Suez consists of a series of low headlands between broad bays, such as Jemsa Bay. Pediments, some wide, gentle and sandy, others steeper and gravelly, decline from arid mountain scarps that angle successively coastward, and end in narrow beaches of sand or shingle, some of which have a few low dunes, bordered by a shallow fringing reef terrace. The pediments continue in front of the North Galala mountain ranges, but are interrupted by a coast range at Gebel Zeit. El Malbaha is an elongated sebkha behind a sandy barrier, and another similar sebkha backs the bay north of Ras Gharib. Coral reefs fringe Ras Ruhami and Ras Azabaraban, and the mountains are close to the coast at Ras Zafarana. A wide valley between steep sandstone escarpments then runs from west to east, with Wadi Araba descending to a coastal lowland, and to the north the plateau has an eastern escarpment dissected by wadis that cross a pediment to the coast. Wadi Ghweibba runs out from another west to east valley to a sebkha that has

This is a revised version of a chapter by Mohamed El-Ashry in The World’s Coastline (1985: 513–517), including contributions by Yaacov Nir.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 15.7.1 Artificial beach at the Holiday Inn, Sarfaga. (Courtesy Geostudies.)

⊡⊡ Fig. 15.7.2 The beach at El Gharib, backed by a broad pediment, with outcrops of Tertiary sandstone in the foreground. (Courtesy Geostudies.)

stunted mangroves, and the coral reefs fade out as the Gulf of Suez narrows northward. The Port of Suez occupies the head of the Gulf at the entrance to the Suez Canal, which crosses an isthmus that emerged from the sea in late Miocene to early Pliocene times to separate the Red Sea from the Mediterranean.

3.  Sinai Peninsula The west coast of the Sinai Peninsula has wadis descending across pediments to marshy sebkhas. Ras Matarama is a large recurved spit sheltering an older, inner spit formation. Sectors of steep coast with some fringing

Egypt, Red Sea Coast

15.7

⊡⊡ Fig. 15.7.3 Dunes behind the beach at El Gharib. (Courtesy Geostudies.)

⊡⊡ Fig. 15.7.4 Beach rock on the shore at Hurghada. (Courtesy Geostudies.)

reefs run south past Ras Sheratib, where a spit partly encloses a bay. To the south, a narrow isthmus links the high peninsula at Ras Muhammad, the southernmost point of Sinai. The coast then swings northeast past Ras Umm Sidd and Sherm el Sheikh, and runs alongside the narrow Strait of Tiran, leading into the Gulf of ‘Aqaba. This is reported

to be the northern limit of mangroves on the Sinai Peninsula. Fringing coral reefs occur intermittently along the Gulf of Aqaba coast, where steep sectors alternate with low forelands and sebkhas at wadi mouths past Dahab to Nuweiba. The coast is then backed by the steep eastern escarpment of the Egma Plateau, with narrow pediments descending to low cliffs, beaches and fringing reefs.

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4.  Israeli Coast, Gulf of Aqaba

References

There is a short sector of coast at the head of the Gulf of Aqaba that is part of Israel. An east-facing steep sector with nearshore coral structures has segments of narrow coastal plain fronting the mountains north from Bir Taba in Egypt. This widens across the border to Eilat in Israel. The coastline curves eastward across the southern end of the Dead Sea Rift Valley, and is fringed by a sandy beach, much modified by tourist development. The coast is crossed by the Jordanian border towards the city of Aqaba.

Gvirtzman G, Buchbinder B, Sneh A, Nir Y, Friedman GM (1977) Morphology of the Red Sea fringing reefs: a result of the erosional pattern of the Last-Glacial low-stand sea level and the following recolonization. Mém Bureau Réch Geol Min 89:480–491 Misak RF, Attia SH (1983) On the sand dunes of the Sinai Peninsula. Egypt J Geol 1–2:115–131 Nir Y (1996) Sediment transport patterns of southern Sinai coasts and their role in the Holocene development of coral reefs and lagoons. J Coast Res 12:70–78 Sneh A, Friedman GM (1984) Spit complexes along the eastern coast of the Gulf of Suez. Sediment Geol 39:211–226

15.8 Jordan

Yaacov Nir

The Hashemite Kingdom of Jordan has only a short (26 km) coastline in the southern part of the country at the head of the Gulf of ‘Aqaba. The Gulf of ‘Aqaba is a part of the East African – Dead Sea – Syrian Rift Valley. The nearshore sea floor declines very steeply, and deltas at the mouths of wadis are consequently very limited. A south-facing Jordanian coastline (about 2.5 km long) extends from the Israeli border east of the modern city of Eilat, and west of the Jordanian city of ‘Aqaba, where it sharply turns southward to the Saudi Arabian border at Ad Durra. ‘Aqaba is a port and coastal resort with a history dating from at least the tenth century bc. The climate of southern Jordan is very warm for about 8 months of the year, when daily temperatures are around 40°C, while winter is mild. Mean monthly temperature at ‘Aqaba region is 16°C in January, rising to 31.5°C in July, and mean annual rainfall is 35 mm. Water entering the Gulf of ‘Aqaba from the Red Sea through the Strait of Tiran undergoes strong evaporation over negligible precipitation, and returns to the Red Sea across the sill partly as an underflow (Emery 1964). As a result of the high temperatures and the absence of perennial streams, salinity of this sector of the gulf is very high, reaching 40.8 ppt. Tides are semidiurnal and uniform along the shores of the Gulf of ‘Aqaba. Winds are mostly from the northern and NW for about 85% of the year, resulting in very small waves. Southerly storms occur mostly during winter, and may generate waves as high as 4 m over to the 180 km long fetch.

The desiccated steep mountain front on the eastern side of the rift valley that is dissected by deep wadis, fronted by a more gently sloping apron of broad, partly confluent sand and gravel alluvial fans. South from ‘Aqaba the mountain front, which is mostly granite, advances to within a few km of the coast. Further south, the terrain is mostly composed of large alluvial ridges 4–10 km wide. It is crossed by numerous watercourses descending from steep-sided wadis, and the depositional lower slope has been truncated to form low cliffs and bluffs along the coast. The rift valley floor ends in a gently curved sandy beach at the head of the Gulf of ‘Aqaba, shaped by southerly wave action. The coast south to the Saudi Arabian border has an almost continuous fringing coral reef, backed by beaches of coralline sand, mixed with quartzose sand, heavy minerals originating in the near-by granite ridges, and gravel from the wadis. Fringing coral reefs are well developed on minor headlands, notably Yamanieh Reef close to the Saudi border. Their outer edge drops to a steeply sloping sea floor.

Reference Emery KO (1964) Sediment of the Gulf of ‘Aqaba (Eilat). In: Miller RL (ed) Papers in marine geology: Shepard commemorative volume. Macmillan, New York, pp 257–273

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_15.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

16.0  South-West Asia – Editorial Introduction

2.  The Southern Coast of Arabia

Arabia occupies a tectonic plate bounded to the west by  the Dead Sea and Red Sea Rifts, to the south by the Indian Ocean and to the east by a zone of faults extending  through Iran to the Gulf of Oman (Sanlaville and Prieur 2005). The Red Sea is a widening rift, considered to be an incipient ocean. It has a narrow, shallow connection with the Indian Ocean (the Rab al Mandab) and because of high evaporation and low rainfall and ­run-off it is more saline than the Indian Ocean, salinity increasing  northward to more than 4.1% in the Gulf of Aqaba. High mountains of Pre-Cambrian and Palaeozoic igneous, volcanic and metamorphic rock rise near the western  margin of the Arabian Plate, parallel to the Red Sea ­coastline, and there is a broad regional slope down eastward to the Persian Gulf, a shallow sea (average depth 35 m) occupying a structural depression. Its eastern (Iranian) coast runs parallel to NW–SE mountain ranges on strongly folded formations, and these swing eastward behind the Makran coast on the northern side of the Gulf of Oman.

The southern coast of Arabia (> Southern Arabia and Oman) is also hot and dry. A narrow lowland is backed by mountains, with steep and cliffed rocky sectors where upland ridges interrupt the coastal plain, as at Ras Fartak, where the cliff is 580 m high. There are recent volcanoes in  the south-west, notably at Aden. The continental shelf is narrow, and a strong southerly swell arrives from the Indian Ocean. The sea is relatively calm during the NE monsoon in winter, but rough in the SW monsoon in summer. There are occasional tropical storms. Mean spring tide range is between 2 and 4 m. The southern coast of the Gulf of Oman has steep mountainous slopes on strongly folded sandstones and limestones in the south and north, trenched by wadis, some with fans extending into the sea. In between is a coastal lowland, the Al Batina Plain, with a gravelly piedmont and a fringe of beaches and dunes. The steep-sided limestone peninsula of Musandam projects into the Strait of Hormuz.

1.  Red Sea Coast

3.  Persian Gulf Coast

The east coast of the Gulf of Aqaba has mountains descending almost to the sea, with only a narrow coastal plain. Similar features extend along the > Saudi Arabia, Red Sea Coast to the south, but the coastal plain widens beyond Yanbu’ al Bahr, and is up to 40 km wide near Jeddah. The mountains are fronted by a gravelly piedmont slope declining to the coastal plain of Quaternary deposits  which ­continues south as the Tihana coastal plain in  southern > Saudi Arabia and >Yemen. The coastal lowlands are terraced with emerged coral reefs, and fringed by  beaches and dunes. The Red Sea coast is hot and arid, cut by bays (sharms), and fringed by intermittent coral reefs, with low-lying saline flats (sebkhas), some with mangroves. Tide ranges are small (around 0.1 m) but winds and pressure variations can cause oscillations of up to 0.9 m. Waves are generated by the prevailing NW winds.

The coast of the > United Arab Emirates is low-lying and bordered by beaches, spits and barriers, with ­extensive supratidal saline flats (sebkhas), up to 10 km wide, some with mangroves. The Dubai coast has been transformed on a huge scale, and is said to have been lengthened from 60 km to almost 1,000 km by the formation of artificial islands and promontories intersected by canals. Beaches and dunes are mainly of biogenic carbonate sand derived from the sea floor, and the nearshore waters are shallow. The mean spring tide range is small (1.3 m at Dubai), and wave action is generated by northerly winds. There are occasional storm surges. To the west, the coast turns northward to the > Qatar Peninsula. Eocene limestones and dolomites form cliffs that are fronted by an emerged zone of sebkhas and coralline tidal flats, and are only occasionally ­submerged by

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_16.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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the sea. The island of > Bahrain and the Damman Peninsula in Saudi Arabia are also of Eocene limestone, with generally low coasts. Long sandy beaches are ­separated by rocky headlands and backed by silty seb­­ khas on the > Saudi Arabia, Persian Gulf to the north and the coast of > Kuwait and Iraq is generally low-lying, with some minor cliffs. Tide ranges increase northward to 3.5 m in the bay of Kuwait, and wave action is generated by easterly winds, particularly in winter. At the head of the Gulf are  low-lying alluvial coasts associated with the Tigris–Euphrates delta, including the mangrove-fringed island of Bubiyan. The eastern (Iranian) coast (> Iran) of the Persian Gulf follows the NW–SE trend of the mountain ranges, which consist of strongly folded limestone and marls of Jurassic, Cretaceous and Lower Tertiary Age, shaped by Alpine earth movements. There are many salt domes, forming islands and shoals. There is a coastal plain of Quaternary sediment up to 40 km wide, with beaches, beach ridges and dunes as well as marshy sebkhas and wide intertidal mudflats (tide ranges between 1 and 4 m). There are cliffs on hilly sectors and around high islands

such as Kharg and Qeshm, and deltas at river mouths, notably of the Hilleh and Mehrar Rivers.

4.  The Makran Coast A narrow coastal plain fringes highlands as the coast curves round the Strait of Hormuz and widens as it continues along the northern side of the Gulf of Oman, the Makran coast. South of the ranges the lowlands include gently folded Pliocene calcareous mudstones and Plio-Pleistocene marine sandstones. There are platforms at several levels, indicating a history of tectonic uplift. The tide range is about 2 m, but exposure to swell from the Indian Ocean increases eastward, and there are bold cliffed headlands.

Reference Sanlaville P, Prieur A (2005) Asia, Middle East, coastal ecology and geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 71–83

16.1 Saudi Arabia, Red Sea Coast

Eric Bird

1.  Introduction The Kingdom of Saudi Arabia has a Red Sea coast about 2,000 km long. The Arabian Peninsula has a steep western escarpment rising to 1,500–3,000 m at the edge of a plateau that declines gradually eastward, flattening to a broad coastal plain. The escarpment crest is only 40 km inland in the north, diverging to 150 km in the south. The landscape is largely desert. Numerous short, steep wadis descend the escarpment to the Red Sea coast, generally dry and gravelly but flowing briefly and torrentially after occasional heavy rain. Longer watercourses run eastward, but there are no permanent rivers. The climate is hot and arid. Jeddah on the Red Sea coast has 23°C in January and 31°C in July, with a mean annual rainfall of 81 mm. The Red Sea has surface temperatures rising to 31°C in the south and 27°C in the north in July, and between 24 and 26°C in January, falling to 20°C in the Gulf of Aqaba. Salinity is high (37–40 ppt) and evaporation from the sea surface of about 2 m/year is compensated by inflow from the Indian Ocean through the Bab el Mandeb. The warm saline water, clear in the absence of river inflow, permits coral growth anywhere in the Red Sea, and coral reefs are extensive, with fringing reefs along much of the Saudi coast, typically with backing shallow lagoons. Tides in the Red Sea are small: Jeddah has a tide range of 0.2 m and Jizan, in the south, 0.8 m, and apart from coral reef flats there is only a narrow intertidal zone. Seas are generally slight (waves less than a metre high) and often calm, but there are occasional storms, particularly in winter, producing waves up to 4 m high. Currents are weak, and mainly wind-driven. The Late Quaternary marine transgression brought the sea to its present level about 6,000 years ago, drowning the mouths of wadis to form the long narrow and deep marine inlets called sharms on the Red Sea coast (Behairy 1983). These are straight or meandering, bordered by coral gardens or reefs, often interrupted at the inner end by sediment washed in from the wadi. Coral reefs that had formed in Pleistocene times and been dissected and reduced during the Last Glacial low sea level phase were recolonised by corals as the marine transgression

proceeded, and the growth of Holocene reefs ensued (Gvirtzman 1977). Patterns of sedimentation have been related to coastal processes by Geith (2000). There was a brief phase of slightly higher sea level followed by a lowering, which resulted in the emergence of some coral reefs and the deposition of beaches and beach ridges behind fringing reefs on the Red Sea coast. Beach rock has formed where beach sand and gravel have been cemented by aragonite. There is evidence of higher Pleistocene coastlines, including emerged coral reefs, on the Red Sea coast, with dislocations resulting from uplift and faulting. The Farsan Islands, off southwest Saudi Arabia, are a group of uplifted reefs with plateaux dislocated by fault scarps, as on Abulat Island (> Fig. 16.1.1).

2.  The Red Sea Coast (North to South) On the eastern coast of the Gulf of Aqaba, mountain slopes descend to a coastline with low cliffs and sandy beaches, with intermittent fringing coral reefs in the southern part. At Ras al Qasbah, the coast becomes more intricate, with headlands, islands and coral reefs, and small bays with sandy beaches. There are occasional reef-edged sharms at the mouths of wadis, as at Sharm Jubbah (27° 30' N). South from Duba, there are fringing reefs, interrupted at the mouths of wadis such as the Wadi Hayyan, which flows down from the Al Hisa mountains and carries sand and gravel to the coast. At Al Wajh, there are some emerged reefs offshore. Reefs and bars run out at Umm Urumah to Mashabih island, on to Shabah and back to the mainland, enclosing a large lagoon. A series of bays past Umm Lajj look out to reefs and some emerged reef islands, and the coast is set forward to Ras Abu Madd. Fringing reefs occur southward, past Ras Baridi and the headlands and bays round to Yanbu al Bahr. South of Yanbu al Bahr, the coastal plain (Tihamat) widens. There is a sharm at Rayyis, and a lagoon has been partly enclosed by the northward growth of a spit at Ras Masturah. Toward Jeddah, the coast is low-lying and reeffringed, with a shallow lagoon backed by a low cliff. The

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⊡⊡ Fig. 16.1.1 Abulat Island, a raised coral island on Farsan Bank, Saudi Arabia. 1 – low coast, 2 – coast with fallen blocks, 3 – overhanging scarp, 4 – double overhang, 5 – successive overhangs, 6 – beach rock, 7 – living coral heads, 8 – fault scarp, 9 – low plain, 10 – karstic plateau, 11 – heights in metres, 12 – emerged coral knoll.

Holocene coral reefs are accompanied by carbonate sediment (Behairy and El-Sayed 1984), and end seaward in slopes plunging to depths of up to 400 m. The low cliff stands at the edge of an apron of downwashed colluviums, which conceals emerged Pleistocene reef terraces. Abhur Creek, north of Jeddah, is a sharm at the mouth of the Wadi Al Korae, bordered by a low cliff with a notch and visor. At Jeddah, the so-called corniche coast has been artificially advanced by dumping rocky debris in the shallow lagoon to make land for urban development and car parks, and the fringing reef shows signs of damage. The coast to the south is set forward to Al Ras al Aswad, and there are a series of bays cut into an emerged coral reef fringe. Offshore, there are a number of emerged coral reef

islands. At Qishran, a lagoon has formed, and several coral reefs run parallel to the coast south of Ras al Askar. They include the Sirraya Reef and Jabbarah Fara. Al Qunfudhah stands on a reef-fringed promontory with sandy bays on either side, and the embayed coast continues south along the Tihamat ash Sham Plain to Am Shuqayq. To the southeast is the Thamat Asir Plain where rivers flowing down from the escarpment pass through a well-vegetated zone, but toward the coast they disappear into the sand. A spit runs southward at Ras at Tarfa, on the low wave energy coast behind the Farsan Islands. The Farsan Islands consist of many reefs and emerged reef islands showing tectonic dislocation (Guilcher 1955). South of Jizan, a long fringing reef extends across the border into Yemen.

Saudi Arabia, Red Sea Coast  

References Behairy AKA (1983) Marine transgressions in the west coast of Saudi Arabia (Red Sea) between mid-Pleistocene and Present. Mar Geol 52:M25–M31 Behairy AKA, El-Sayed MK (1984) Carbonate sediment in a modern sea reef north of Jeddah, Saudi Arabia. Mar Geol 58:443–450

   16.1

Geith AM (2000) Sedimentary reflection of the coastal processes on the shoreline sediment along the eastern Red Sea coast, Saudi Arabia. Z Geomorphol 44:449–468 Guilcher A (1955) Geomorphology of the northern part of the Farsan Bank (in French). Ann Inst Océanogr 30:55–100 Gvirtzman G (1977) Morphology of the Red Sea fringing reefs, Vol 89. Memoirs, Bureau de Recherche Géologiques et Minière, Paris, pp 480–491

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16.2 Yemen

Eric Bird

The coastal landforms of the Republic of Yemen in ­southern Arabia have received very little attention from geomorphologists, but are shown on the US Geological Survey map of the Arabian Peninsula, 1–270A (1963) at 1:2 million. The climate is arid, and waves are generated across the narrow Red Sea, mainly from the SW. Tide ranges are small, generally about a metre. The Red Sea coast of Yemen is a continuation of the Tihamah coastal plain of southwest Saudi Arabia, an irregular arid coast with sandy beaches and fringing coral reefs and a group of islands including Kamaran. There are also several islands and irregular promontories of denuded volcanic rock (the Aden Volcanics, of Late Cainozoic age). Sebkhas with saline crusts are extensive, and barchans 4–14 m high drift across them. Near the coast some of the dunes become vegetated on the windward slopes, and may evolve into parabolic dunes. Salt domes form hills, including the Al Sarif peninsula, which is a tombolo. Sand drifting northward has built the Kathib spit near Al Hudaydah

(Youseff 1991). South from Al Hudaydah the wide coastal plain fronts a steep, dissected arid escarpment, and along its seaward margin are low cliffs and beaches. Coral reefs occur intermittently along this coastline, and there are mangroves and salt marshes in sheltered areas and nearshore seagrass communities (Zahran and Al Kaf 1996). The Red Sea narrows to the Bab al Mandab Strait, which is bordered by a steep coast with cliffs on the hilly peninsula of At Trubah. About 800 km to the east of Aden, in the Gulf of Aden, is the Yemen island of Socotra (> Indian Ocean Islands).

References Youssef EA (1991) A note on the geomorphology of the coastal plain between Al-Hudaydah and Alsalif Peninsula, Red Sea coast, Yemen Arab Republic. Geogr J 157:71–73 Zahran MA, Al Kaf HF (1996) Introduction to the littoral halophytes of Yemen. Arab Gulf J Sci Res 14:691–703

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_16.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

16.3 Southern Arabia and Oman

Eric Bird

1.  Introduction The coastal landforms of the Sultanate of Oman in southern Arabia have also received very little attention from geomorphologists, but are shown on the U.S. Geological Survey map of the Arabian Peninsula, 1–270A (1963) at 1:2 million and described in Pilot Handbooks 63 and 64 published by the Hydrographer of the British Navy.

2.  Southern Arabia Cape Bab-al-Mandab marks the beginning of the southfacing Arabian coastline, which extends past Aden eastward to Oman. The coastline receives Indian Ocean swell arriving from a southerly direction, accompanied in ­summer by waves generated by SE trade winds. It is a hot and arid coast (mean annual rainfall at Aden is 38 mm; at Salalah 112 mm; at Masirah Island 38 mm; at Muscat 100 mm) backed by desert landscapes with at best a sparse, shrubby vegetation. However, monsoonal rains in the Yemen uplands produce a number of streams draining wadis and flowing into deltaic fans on sectors of coastal plain in the southwest, and the Samham Hills in SW Oman also receive sufficient monsoonal rainfall to sustain a relatively rich vegetation. Mean spring tide ranges are moderate, generally between 2 and 4 m: Aden has 1.5 m and Al Mukullah 1.2 m. At the entrance to the Red Sea are the large tombolos of the Aden district. Longshore drifting to the west by waves generated by SE trade winds has built the spit that shelters the bay lagoon at Khawr am Umayrah, west of Aden, in which broad tidal mudflats are exposed at low tide. The hilly volcanic promontory of Jabal Ibsan (Little Aden) has small bays between headlands such as Ras al Araja, and is separated from Jabal Shamsan (the Aden Peninsula) by the bay of Bandar Tawahi, backed by a curving sandy beach that has been shaped by refracted southerly swell. Aden stands on the rim of a dissected and partly ­submerged extinct volcano, the coastal cliffs being cut into lavas, agglomerates, and ash, with basalt dykes prominent. Long sandy beaches are backed by dunes on

a coastal plain up to 20 km wide, behind which slopes rise to an interior plateau. Raised beaches and emerged coral reefs, horizontal or slightly tilted, occur at intervals along the Gulf of Aden coast (Beydoun 1960). Some wadi fans end in lobate deltas, as at Abyan (Wadi Bana outwash) and Wadi Ahwar, and occasional flooding after heavy rains in the uplands delivers sandy sediment to beaches. At Shuqra, the coastal plain is narrow and gravelly in front of the volcanic Khaur Mountains, and carries a string of small volcanic cones. Ras al Ratl is a high, round volcanic promontory, and Karif Shawrab has a brackish, ­mangrove-fringed crater lake. Valleys opening to the coastal plain include Wadi Mayfa’ah and Wadi Hajr, the only perennial stream. To the east, upland spurs reaching the coast end in high cliffs of Tertiary limestone, as at Ras al Mukalla, and there are more coastal lava flows and volcanic cones between Shihr and Sayhut. Locally, the black lava has flowed down wadis incised into white limestone to outcrop in cliffs. The largest river valley in southern Arabia, the Wadi Masilah, draining from the Hadhramaut Trough, crosses the coastal plain here, but strong ocean swell has prevented the building of a protruding delta. Mangroves occur in swampy sectors, but surf-washed beaches of white sand and occasional gravel line the coast, which becomes cliffed at Sayhut. Dune sands dominate coastal plains east of here, until the coastline changes orientation at the bold rocky headland of Ras Fartak, where a steep sector with vertical cliffs up to 580 m high truncates the east–west structures of the Mahrah Highlands south of Tabut on the shores of Qamar Bay. Behind Qamar Bay, the coastal plain widens again, consisting of convergent braided wadi outwash fans with dunes along the seaward margin.

3.  The Oman Coast The coastal plain narrows across the Oman border at Ras Darbat toward the small fishing port of Salalah. The Oman coast is very dry (annual rainfall less than 100 mm), with hilly topography in the Dhufar hinterland (Tschopp 1967).

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It is exposed to Indian Ocean swell arriving from the south, accompanied and waves generated by SE trade winds. The tide range is generally about a metre, increasing to 1.9 m at Muscat. The coastline has a sequence of large bays with sandy beaches washed by ocean swell separated by steep rocky promontories. Pre-Cambrian metamorphic rocks form the lowland fronting the Samham Hills, which are bordered by deep sea close inshore, but eastward the continental shelf widens past Swaquirah Bay and Masirah Island toward Ras al Hadd. The five rugged Kuria Muria Islands offshore consist of granite and volcanic conglomerate, with occasional gravel beaches interrupting a rocky coastline. Behind Swaquirah Bay, the coastal plain widens and coral reefs surround the shore. The coastline turns northward past Ras Madhraka, a promontory of metamorphic rock, and Palaeozoic and Pre-Cambrian basement outcrops along the western side of Masirah Bay, where again there are coral reefs. The sand ridges of the Wahibi desert run north–south and reach the coast in the lee of Masirah Island, where shoals and reefs reduce wave action so that the Hikman Peninsula has a broad sabkha fringe, flat sandy terrain inundated occasionally by the sea. Masirah Island, parallel to the coastline, consists of igneous and metamorphic rocks and has generally steep coasts bordering rugged uplands: here, as elsewhere in southern Arabia, these rocks weather to jagged crests and rubbly slopes, in contrast with the more tabular or scarped limestone topography.

Northeastward, sandy surf beaches line the coast as far as the low sandy promontory of Ras al Hadd, where the coastline swings NW alongside the Gulf of Oman. Wave energy diminishes as ocean swell is refracted round Ras al Hadd, and the southern shores of the Gulf of Oman receive waves generated mainly by the southeast trade winds. Just west of Ras al Hadd are low limestone cliffs, interrupted by small tidal inlets, and by a larger sharm (Wadi Falaij) at Sur. Between Sur and Muscat, the coastal slopes show terraces marking emerged coastlines at various levels up to 250 m above sea level at the mouth of Wadi Ali. In the hinterland, the bleak arid ranges of the Hajar Mountains (Jabal Al Akhdar) attain more than 3,000 m. The coastline runs parallel to the strike of igneous and metamorphic rocks, flanked by dipping Cretaceous limestones, which form these highlands, their seaward slopes incised by many wadis that open to a narrow coastal plain. Muscat is sited on a cliffed coast cut in weathered ­serpentine. A coast road, protected by banks of concrete tetrapods, has been built in front of the cliffed coastline (>Fig. 16.3.1). To the west, past the oil terminals of Matrah, the Batina coastal plain develops and widens to 16 km. A fertile lowland with acacia trees, date palms, and irrigated plots around spring-fed oases, it is bordered by sandy beaches, locally becoming spits built by northwestward longshore drift to enclose minor lagoons and salt marshes. Northward the foothills approach the coast, and there are marine terraces, including emerged coral reefs, about 10 and 28 m above present sea level. ⊡⊡ Fig. 16.3.1 Artificial coast resulting from highway construction near Muscat. (Courtesy Embassy of the Sultanate of Oman.)

Southern Arabia and Oman  

16.3

⊡⊡ Fig. 16.3.2 Southward-dipping limestones and shales influence the outlines of the Musandam cliffs. (Courtesy Philippa Barr.)

⊡⊡ Fig. 16.3.3 Coastal terrace fronting mountains on the Musandam coast, limestone cliffs with a basal notch. (Courtesy Philippa Barr.)

The south-eastern coast of the Sultanate of Oman, known as Dhofar, has a climate with a strong seasonal contrast between the very dry hot season (NE Monsoon), when shore features are desiccated, and the SW Monsoon (Khareef) in July to September, when winds blowing along the coast cause upwelling of deep cold oceanic water and produce cool, misty weather and rapid vegetation ­recovery in coastal regions. Shore environments are then colonized by algae, kelp, Ulva, Sargassopus, as well as

rock oysters and barnacles. There are also coral gardens, solid reef building being impeded by the cold water season. As the peninsula narrows toward the Strait of Hormuz (past a sector of coastline that belongs to the United Arab Emirates), ridges of Jurassic and Cretaceous limestone culminate in the highly indented Musandam Peninsula (>Figs. 16.3.2 and >16.3.3), an area of gently folded and much-faulted limestone with steep coasts bordering

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⊡⊡ Fig. 16.3.4 Limestone cliffs beside a ria on the Musandam coast. (Courtesy Philippa Barr.)

branching rias formed by the marine submergence of wadis (>Fig.  16.3.4). There are no raised beaches here, and it is thought that submergence by the Holocene marine transgression was augmented by downwarping of the land (Vita-Finzi 1973). Cliffs have formed where the steep coastal slopes are being undercut by the sea, especially where shales outcrop beneath the limestone. Rockfalls and landslides are common. Many of the inlets have sand or pebble beaches at the mouths of wadis, where cemented gravel deposits also form low coastal benches.

Locally, there are coral reefs. Out beyond Ras Musandam, there are towering limestone stacks, the Quoin Islands.

References Beydoun ZR (1960) Synopsis of the geology of East Aden Protectorate. Proc 21st Int Geol Congre 21:131–149 Tschopp RH (1967) The general geology of Oman. Proc of the 7th World Petrol Congre 2:231–241 Vita-Finzi C (1973) The Musandam expedition. Geogr J 139:414–421

16.4 United Arab Emirates

Eric Bird

1.  Introduction The United Arab Emirates have a coastline of about 800 km on the southern shore of the Persian Gulf, extending from the base of the Qatar Peninsula to the mountainous Musandam Peninsula (Drew 1985). The length of the coastline has increased as the result of artificial development, particularly in Dubai. The coastline has also increased where shoals offshore have been reclaimed to form islands that simulate The World (Hansen 2005). The climate is hot and dry, and the landscape mainly desert. Dubai has a temperature of 23°C in January and 42°C in July, and a mean annual rainfall of 60 mm. Tides are semidiurnal, with a mean spring tide range of 0.9 m at Dubai, increasing both eastward to 1.8 m at the Musandam Peninsula and westward to about 1.5 m in Qatar. The prevailing winds are northwesterly and strong during the shamal, which blows off the northern Arabian Desert. In summer the northwest winds blow almost constantly, whereas in winter they are generally stronger but of shorter duration. Calm or slight seas (wave height up to 0.9 m) occur more than 75% of the time in the Persian Gulf, and rough, high seas (wave height about 1.5 m) 5–6% of the time, usually in autumn and winter. Water flows into the Persian Gulf through the Strait of Hormuz, up along the northern shore, and down along the southern shore, resulting in an eastward flow off the United Arab Emirates. Salinity is high because of strong evaporation and the lack of rainfall. In Abu Dhabi, for example, evaporation rates average 25 mm/day, with a maximum of 55 mm/day, while rainfall averages only about 40 mm/year. Salinity varies through the year, according to evaporation rates and current circulation; minimum salinities occur in August, averaging 39.0–39.5‰, but in protected bays and lagoons (such as the inner lagoon of Abu Dhabi) salinity can be over 6% (Schneider 1975). In the absence of rivers the terrigenous sediment supplied to the coast from the Arabian land mass consists of wind-blown fine sand and dust (Sugden 1963). The southern part of the Persian Gulf is a region of very high ­biogenic sediment production, with many coral and algal reefs. Coastal and sea floor sediment are dominated by calcium

carbonate precipitated from sea water and the remains of  shelly organisms, corals and algae. Coastal sediment are  distributed by waves and tidal currents in patterns related to the orientation of the coast, the protection of the  Qatar Peninsula which acts as a windward barrier, and the Great Pearl Bank, an offshore reef and shoal complex (Purser and Evans 1973). Physical characteristics of beaches are discussed by Alsharan and El-Sammak (2004). The western and central parts of the coast are dominated by the presence of sebkhas. These are low-lying saline flats that stand just above normal high tide level, and are submerged at the highest tides and during storm surges: they are lagoon floors that have partly emerged. They extend for over 300 km westward from Ras Ghanada, and have a maximum width of about 30 km. They are bordered on the landward side by low bluffs cut in Tertiary and Quaternary formations (Evans et al. 1964). Sebkhas are floored with unconsolidated carbonate sediment and minor amounts of quartz and other minerals, including evaporites. A series of low linear ridges runs parallel to the coast, through the sebkha, especially in the western region. These old beach ridges decrease in height inland and, like the present seaward ridge, are composed largely of gastropod shells. In the central region these ridges are replaced by old strandlines and transverse tidal drainage channels (Evans et al. 1964). Mangroves grow locally in sheltered areas, notably on the shores of lagoons. The sebkhas in the vicinity of Abu Dhabi have a basement of Miocene and Pliocene carbonate rocks, overlain by 4–5 m of unconsolidated sediment, typically consisting of aeolian, cross-bedded quartz and carbonate sands and fossiliferous carbonate muds, interbedded with sediment lenses up to 3 m thick. Sebkha development in the central region began more than 7,000 years ago, when late Pleistocene and Holocene sediment along the coast were dominated by aeolian sands. The Late Quaternary marine transgression culminated about 4,000 years ago when the sea rose 0.5 m above its present level, and there was extensive carbonate sedimentation. An ensuing emergence produced a wide supratidal and intertidal lagoon zone and formed broad dune-capped

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_16.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 16.4.1 Barrier islands, lagoons and sebkhas near Abu Dhabi. (Courtesy Geostudies.)

barrier islands, as well as algal and coral reefs and nearshore shoals (>Fig. 16.4.1) (Evans et al. 1969). The dunes on the barrier islands are bare of vegetation and broad and irregular in form: had there been a vegetation cover they may have become multiple foredunes. Parabolic dunes may be related to salt crusting, retaining parts of the sand surface. Deposition of carbonate sand and mud, growth of calcareous organisms, and accretion of wind-blown dust from inland and offshore islands continue in the lagoons and the intertidal and supratidal environments. The sebkhas are also sites of formation of evaporite minerals, gypsum and oolites, and eventually dry out as saline flats. Progradation has been intermittent, with stages indicated by parallel beach ridges on old strandlines. Some beach ridges have been built during storms, particularly in the eastern and western sebkhas, the central zone being sheltered by reefs, shoals and barriers. Others formed as spits extending from low headlands and knolls of older rock, with subsequent infilling of landward bays. Sebkhas may be arenaceous or argillaceous. Those forming now are generally arenaceous, caused by aeolian sand filling a lagoon or embayment and being reworked by waves and currents. Their surfaces are flat and smooth, with gradients of less than 0.5 m/km. Water rising through pore spaces in the sand coats the grains and increases lubrication, so that these sebkhas develop into soft quicksand with low bearing strength. Argillaceous sebkhas, also numerous along the coast, were formed from the production of calcareous mud by algae and other organisms in shallow coastal embayments. This wet, soft and sticky mud has a very low bearing strength, and flows under pressure. Aeolian sand blowing across the surface adheres to this

mud, and forms a layer up to 1 m thick; when not covered by sand, the mud dries and hardens in the summer heat. The outer coast of the barrier islands consists of unvegetated dunes behind broad gently shelving beaches of brown calcareous oolitic sand descending to shallow sandy nearshore areas, and at low tide broad sandy flats are exposed, threaded by deep channels. Beach rock occurs where beach sediment have been cemented by aragonite (Shinn 1969). Coral reefs are extensive in coastal waters, and there are large areas of seagrass growing on the sandy sea floor. Outlying islands include several salt domes, notably the island of Das, formed by diapric uplift around an expanding mass of salt (>Fig.  16.4.2). They are bordered by beaches and trailing spits, sometimes with fringing coral reefs on the windward shore.

2.  The United Arab Emirates Coastline The United Arab Emirates include a sector of east-facing coast facing the Gulf of Oman, on the eastern side of the high barren limestone mountain chain that runs north to the Musandam Peninsula. The coast is steep and rocky, with valley-mouth bays and inlets as at Kalba and Khawr Fakkan between rugged headlands. The northeastern region of the United Arab Emirates has the limestone mountains of the Musandam Peninsula, which pass inland behind a sandy low coast extending to Ras Ghanada. Storm-built beaches are backed by coastal dunes of carbonate sand, and several subparallel carbonate sand spits 5–10 km long have grown alongshore, curving around to the northeast. Longshore spits form barriers that extend laterally near Umm al Qawain for more

⊡⊡ Fig. 16.4.2 Dalma Island, a salt dome island in the southern part of the Persian Gulf. (Courtesy Geostudies.)

United Arab Emirates

than  15 km (Bernier et al. 1995). These barriers enclose lagoons dominated by carbonate muds and dissected by channels with associated ebb-tidal deltas; in places there are dry saline flats. The coastline faces northwest and the onshore shamal winds generate strong waves across a long fetch to break on a steep nearshore slope, building beaches and moving sediment northeast along the shore. Some of the lagoons have been filled with calcareous sediment to form sebkhas, across which dunes have migrated (Purser and Evans 1973). Dubai stands on either side of a tidal inlet, Dubai Creek, which has been maintained by recurrent dredging. Development has rendered much of the coastline artificial, but with emplaced sandy beaches (>Fig.  16.4.3). Protruding from the natural coastline is Palm Jumeirah, which has eight branches on either side of a spur projecting from the coastline and sheltered by curved platforms (>Fig. 16.4.4). Two more such structures (Palm Jebel Ali and Palm Deira) are under construction. The central coast, from to Ras Ghanada to Jebel Dhanna, is strongly influenced by the Great Pearl Bank, a chain of algal and coral reefs and shoals. This offshore bar complex runs obliquely to the coast, parallel to the structural axis of the Persian Gulf, and part of the late Tertiary tectonic system (Kassler 1973). Its eastern part is shallow, a broad ridge surmounted by a series of carbonate sand shoals and small islands, but to the west it diverges and deepens to between 5 and 10 m, ending about 50 km offshore from Jebel Dhanna. Wave energy on the coastal barriers is diminished by the protection afforded the coast (Purser and Evans 1973). The barrier islands of the central region show gradational changes in form along the coast. The barrier islands northeast of Abu Dhabi have spits growing at their eastern ⊡⊡ Fig. 16.4.3 The beach at Dubai. (Courtesy Philippa Barr.)

16.4

⊡⊡ Fig. 16.4.4 Palm Jubeirah, a protrusion from the Dubai coastline. (Courtesy Philippa Barr.)

ends, because of longshore transport due to the long fetch across the Gulf (Purser and Evans 1973; Schneider 1975). As the fetch diminishes, a westward counter-drift becomes effective, and the barrier islands have spit accretion at both ends. Further to the west the barrier islands develop trailing spits at right angles to the barrier axis, occasionally forming tombolos. They shelter wide intertidal flats composed of pelletal sands with abundant gastropods and well-developed blue-green algal mats. The inner coastline is fringed with small beach ridges of gastropod sand and coralline gravel, or small mangrove swamps. Supratidal sediment landward of these coasts contain dolomite and evaporite minerals. The lagoon between the Great Pearl Bank and the coast widens and deepens westward, reaching a width of 140 km and a depth of 40 m at its western end. Salinity of the lagoon water diminishes westward from 60 to 40‰ (Purser and Evans 1973). Yas Island is a salt dome island close to the coast, separated from the Jebel Dhana salt dome promontory by a shallow strait with tidal channels. To the west is an open embayment, backed by the Sebkha Matti. A high (1–3 m) storm beach extends 55 km west to Jebel Baraka. The eastern part consists of red algal debris derived from adjacent fringing reefs, the central part of molluscan shell sand and gravel, and at the western end are oolitic sands derived from adjacent tidal flats. The storm beach is bordered seaward by a sheet of carbonate sands (Purser and Evans 1973). The western coast is protected from the full force of wind-driven waves and currents by the Qatar Peninsula. It is made up of a series of north–south oriented rocky

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promontories separated by embayments that are probably delimited by tectonic fractures (Kassler 1973). The ­promontories have fringing reefs, despite the high salinity (45–50‰). The embayments are about 5 m deep, and their shores are backed by wide intertidal flats and narrow beaches (Purser and Evans 1973). The beach faces, better developed on the more exposed sectors, are usually covered with deposits of driftwood, and the beach ridges are often capped by dunes with sparse vegetation (Evans et al. 1964). Landward of these beaches are sebkhas up to 3 km wide, passing westward into the low-lying isthmus that attaches Qatar to the Arabian mainland.

References Alsharan AS, El-Sammak AA (2004) Grain-size analysis and characterization of sedimentary environments of the United Arab Emirates coastal area. J Coast Res 20:464–477

Bernier P, Dalongeville R, Dupuis B et al (1995) Holocene shoreline variations in the Arabian Gulf: example of the Umm al-Qowayn Lagoon (UAE). Quaternary Int 29–30:95–103 Drew KS (1985) United Arab Emirates. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 723–727 Evans G, Kendall CGSt.C, Skipwith P (1964) Origin of the coastal flats, The Sabkha, of the Trucial Coast, Persian Gulf. Nature 202:759–761 Evans G, Schmidt V, Bush P, Nelson H (1969) Stratigraphy and geologic history of the sabkha, Abu Dhabi, Persian Gulf. Sedimentology 12:145–159 Hansen B (2005) Artificial islands reshape Dubai coast. Civil Eng 75:12–13 Kassler P (1973) The structural and geomorphic evolution of the Persian Gulf. In: Purser BH (ed) The Persian Gulf. Springer Verlag, Berlin, pp 11–32 Purser BH, Evans G (1973) Regional sedimentation along the Trucial Coast, Persian Gulf. In: Purser BH (ed) The Persian Gulf. Springer Verlag, Berlin, pp 211–231 Schneider JF (1975) Recent tidal deposits, Abu Dhabi, UAE, Arabian Gulf. In: Ginsberg RN (ed) Tidal deposits. Springer Verlag, Berlin, pp 209–214 Shinn EA (1969) Submarine lithification of Holocene carbonate sediment in the Persian Gulf. Sedimentology 12:109–144 Sugden W (1963) Some aspects of sedimentation in the Persian Gulf. J Sediment Petrol 33:355–364

16.5 Qatar

Eric Bird

1.  Introduction The Qatar Peninsula interrupts the east coast of Saudi Arabia in the southern part of the Persian Gulf (> Fig. 16.5.1). It has a coastline of about 550 km, and was formerly an island, connected to the mainland by the Holocene emergence of a low-lying saline isthmus. The climate is hot and dry. Doha has a temperature of 17°C in January and 37°C in July, and a mean annual rainfall of 62 mm. Tides are semidiurnal, with a range of about 1.2 m at Masay’ia and 1.5 m at Doha. The beaches and rocky shores are gently shelving to shallow sandy nearshore areas, and at low tide broad tidal flats are exposed. The prevailing winds are northwesterly and strong during the shamal, which blows off the northern Arabian Desert. On the east coast, waves generated by northerly winds produce longshore drifting to the south. Seas are generally slight (waves less than 1 m high) and often calm, but there ⊡⊡ Fig. 16.5.1 Coastal geomorphology of Qatar. (Courtesy Geostudies.) Ras Tanura

Qatif

coral reef

Dammam

marsh, swamp

Manama

sand Bahrain

Halul

Hawar Gul

Dukhan

f of Salw a

SAUDI ARABIA

Doha

QATAR Dalma

are occasional storms, particularly in winter, producing waves up to 4 m high. Currents are weak, and mainly wind-driven. Salinity is high because of strong ­evaporation, the sea surface having 39–41‰, increasing to 41–43‰ at  greater depth. It is higher (55–70‰) in the south of the  Salwah Gulf, and may exceed 70‰ in some coastal lagoons. The Qatar Peninsula is anticlinal, cut in Tertiary formations dominated by Eocene limestones, dolomites and chalk. In the south the coast is generally low-lying, with extensive sandy plains, intertidal hypersaline areas (sebkhas) and coastal lagoons, while the north coast is cliffed and rocky with emerged beaches of Last Interglacial age, including rounded beach gravels 3.3 m above sea level on benches cut in Eocene limestone. The presence of dune calcarenites along coasts where there are now no beaches is probably the outcome of sea level oscillations. The Late Quaternary marine transgression attained the present sea level about 6,000 years ago, and there was a  slightly higher sea level stand in Holocene times. The ensuing emergence left a former cliffed coastline behind a wide zone of beach ridges, dunes and intertidal flats, particularly in the south. The island is generally low-lying, with an anticlinal ridge attaining 98 m on the northwest coast near Dukhan. The land is sparsely vegetated, with much stony and sandy ground. There is much wind-blown sand and dust (Sugden 1963), and dunes including barchans move from NW to SE across the southern part of the island, with sand spilling on to the shore south of Umm Said (Meigs 1966). Rocky coasts are limited, and beaches of brown sand line much of the coast. Fluvial sediment supply is negligible, and apart from the local supply of wind-blown quartzose sand from spilling dunes the beaches are dominated by carbonate sand and gravel, largely derived from coral and algal reefs and shore formations and shelly organisms (Purser 1973). Beach rock is seen where beach sediment have been cemented by aragonite. Coral reefs are extensive in coastal waters, and there are large areas of seagrass growing on the sandy sea floor. Mangroves grow locally in inlets and sheltered areas, notably on the edges of lagoons.

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Narrow inlets, termed khawr, are similar to the sharms of the Red Sea, as at Khawr Al Udeid. There are numerous small Ras, which are cliffy limestone spurs, as at Ras Laffan, Ras Abu Aboud, Ras Rahan and Ras Ashrij. Coastal sebkhas are often partly enclosed by spits or beach ridge barriers. They may originate as emerged lagoons which are at first periodically submerged by the sea, then filled with wind-blown and inwashed sediment and precipitated material including gypsum and oolites, eventually drying out as saline flats. Qatar has a number of outlying islands, notably Halul, site of an oil complex, which is a rugged limestone island with dunes and calcareous beaches. There are also sand shoals and salt dome islands with a central area of upheaved rocks, as on Daz and Zirkuh islands to the east.

2.  The Qatar Coastline On the southeast coast of Qatar is a large lagoon and sebkha association, bordered seaward by spits. The hypersaline Khawr al Udeid lagoon is fringed by dunes on emerged Holocene sebkhas, and has deposits of gypsum and oolitic sand (aragonite-coated grains). To the north are sandy beaches backed by the dunes of the Nijian Qatar desert. There are dune ridges spilling into the sea south of Umm Said, particularly when the strong shamal wind is blowing from the NW, nourishing beaches with desert sand. The beaches have been prograding at 1–2 m/year noses of spilling sand. To the north there are elongated

sand ridges parallel to the coastline, indicating emerged coastlines. Musay’id is an oil tanker port, and north of Umm Said a spit shelters a small harbour in the Bay of Al Bushairiya. There are coral reefs and shoals offshore, and these converge northward to form a sanded fringing reef along the coast. Longshore spits enclose shallow lagoons (> Fig. 16.5.2), as in the Al Wakrah embayment, where a spit built by southward longshore drifting curves southwest. At Ras Abu Aboud, a reef-fringed headland, the coast swings west to Doha, where the urban seafront faces north across a shallow bay. The coastline has advanced as the result of land reclamation here. The coast is embayed north to Al Khawr (Khor), where longshore spits and small barriers with beach ridges stand in front of lagoons and low cliffs cut in Eocene dolomite. The North Khor and Dakhirah lagoons are bordered by sebkhas submerged only at the highest tides. Elongated ridges of carbonate sand have been driven landward, and spits are recurved at the southern ends. This is a good area for the study of landforms built by wind and wave action in the absence of  land vegetation. There is exchange of sand between beaches and nearshore sand bars along the northeast coast of Qatar (> Fig. 16.5.3). The coast swings northwest at Ras Laffan, past Fawayrit to Ras Rakan, the northernmost point of Qatar, where reefs and algal sands extend seaward. The northwest coast is embayed, with low cliffs and segments of emerged beach. A deposit of shelly sand 1.5 m above high tide level was dated at 5,378 ± 80 years bp. The northwest coast is ⊡⊡ Fig. 16.5.2 Paired spits border a lagoon on the northeast coast of Qatar. (Courtesy Geostudies.)

Qatar

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⊡⊡ Fig. 16.5.3 Beach and sand bars on the northeast coast of Qatar. (Courtesy Geostudies.)

bordered by a coral reef, separated from the mainland by a broad moat. At Dukhan the coast is rocky, facing out to the Hawar archipelago. To the south, the coast fringing the shallow Salwah Gulf has high salinity (55–70‰) and the sea floor has quartzose desert sands overlain by marine carbonate sand deposited during the past 7,000 years.

References Meigs P (1966) Geography of coastal deserts, Arid Zone Research, 28. UNESCO, Paris Purser BH (ed) (1973) The Persian Gulf. Springer-Verlag, Berlin Sugden W (1963) Some aspects of sedimentation in the Persian Gulf. J Sediment Petrol 33:355–364

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16.6 Bahrein

Eric Bird

1.  Introduction The island of Bahrein has a coastline 126 km long, and is accompanied by more than 30 small islands. Bahrein is linked to Saudi Arabia by a causeway. The island is lowlying and rocky, a dome structure with a central core of Eocene limestones, sandstones, and marls, the land rising to Jabal Dukhan, an interior hill with a summit 135 m above sea level (Willis 1967), surrounded by in-facing escarpments. There is a broad coastal fringe of unconsolidated Quaternary silts, quartzose dune sands, carbonate sands in beaches and beach ridges, shell deposits, and evaporites. The southern half of the island is a barren sandy plain with some salt marshes, and the northern half is a broad plateau with settlements and gardens irrigated from artesian wells. Groundwater seeping from Saudi Arabia is under sufficient pressure to generate submarine freshwater springs off the north coast. The climate is hot and humid. Bahrein has 19° in January and 36° in July, and the mean annual rainfall is 130 mm. Tides are semidiurnal, with a mean spring tide range of 1.8 m at Mina Salman. The prevailing winds are northwesterly, strong during the shamal, which blows off the northern Arabian desert, and occasionally a hot dry northerly, the gaws, brings sand and dust. Waves generated by the northerly winds produce longshore drifting to the south on both the east and west coasts of Bahrein. Seas are generally slight (waves less than a metre high) and often calm, but there are occasional storms, particularly in winter, producing waves up to 4 m high. Currents are weak, and mainly wind-driven. Salinity is high because of strong evaporation, the sea surface having about 40%0. It is higher (55–70%0) in the south of the Salwah Gulf, and may exceed 70%0 in some coastal lagoons. The Late Quaternary marine transgression brought the sea to its present level about 6,000 years ago (Doornkamp et al. 1980), and there was a brief phase of slightly higher sea level followed by emergence, which resulted in the deposition of beaches, beach ridges, and spits in front of a low cliffed coast, enclosing a number of small lagoons and intertidal sebkhas linked to the sea through narrow inlets. The beaches and rocky shores

shelve gently to shallow sandy nearshore areas, and at low tide broad tidal flats are exposed. Beach rock has formed where beach sand and gravel have been cemented by ­aragonite (Shinn 1969). There are coral reefs, particularly off the north and northeast coasts, and the sea floor is carpeted with bioclastic sandy sediment resting on an extensive sea floor rock pavement produced by Holocene submarine lithification (Shinn 1969). There are large areas of seagrass growing on the sandy sea floor. Narrow inlets, termed khawr, are similar to the sharms of the Red Sea, as at Khawr Al Udeid. There are numerous small Ras, which are cliffy limestone spurs, as at Ras Laffan, Ras Abu Aboud, Ras Rahan, and Ras Ashrij. Coastal sebkhas are found in the south and southwest of the island. They are often partly enclosed by spits or beach ridge barriers. They may originate as emerged lagoons, which are at first periodically submerged by the sea, then filled with wind-blown and inwashed sediment and precipitated material including gypsum and oolites, eventually drying out as saline flats. Mangroves grow locally in inlets and sheltered areas, notably on the edges of lagoons and sebkhas.

2.  The Coastline of Bahrein The northeast coast of Bahrein has been much modified by reclamation, particularly around Manama and the port area. The urbanised island of Al-Muharraq is linked to the main island by a causeway, and surrounded by wide sandy shoals. A small valley opens into the Khawr al-Kab embayment, which is sheltered by the island of Sitrah on the southern side. To the south, the coast becomes embayed, with low promontories at Ras Salbah, Ras Zuwayyid, Ras Abu Jarjur, and Ras Hayyan, which is on a SW-NE syncline. There are wide intervening sandy bays. Northerly winds generate waves that cause southward drifting, forming spits that flank intertidal sebkhas along the coast past Ad Dur, where another SW-NE syncline crosses the coast. A broad tidal flat is exposed at low tide, with reef patches and numerous sand bars, the unconsolidated sand resting on

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_16.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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thin sea floor layers of sandstone cemented by calcium carbonate. At the southern end of the island, sand that has drifted  down the east and west coasts has accumulated in  a series  of sub-parallel beach ridges. These converge southward into an elongated straight spit, the Ras al Barr (>Fig. 16.6.1). The sand is accompanied by broken fragments of thin sandstone and beach rock, particularly on the western side. The spit has been shaped by converging NW and NE waves, and a strong tidal current flows around the end. It continues as a submarine shoal southward almost to the Saudi Arabian coast. The west coast of Bahrein is mainly low and sandy, with longshore spits that have grown southward, and several small lagoons and sebkhas, such as the Dahwat alMumattalah. The beaches shelve very gently, and a wide area of sand bars and some reefs is exposed as the tide falls. There are a number of small headlands with low cliffs. The shallow sea between Bahrein and Saudi Arabia has elongated shoals and channels shaped by tidal currents and the movements of water generated by the NW and N winds, particularly during the shamal and the gaws. The northern coast of Bahrein is bordered by low cliffs and bluffs, with sandy beaches in wide embayments. There are many outcrops of beach rock, and coralline sand and gravel has been derived from nearshore and fringing reefs.

The sea floor has been dredged for sand and gravel used in coastal reclamation, particularly at Manama. Beach rock was used in the construction of the Portuguese fort on the northwest coast. There are several outlying islands and reefs. To the east is Jaradah Reef, a patch reef with an elongated sand ridge (cay) on its southeastern shore, culminating in a spit at the southern end, marked by an outlying tower, The beach at this point slopes steeply below low tide level, and is littered with broken beach rock. Hawar Island, across the Gulf of Bahrein to the southeast, is low-lying and barren, fringed by numerous sandy beach ridges (>Fig. 16.6.2) in patterns that show progradation interrupted by several phases of erosional truncation. There are numerous sand bars in the surrounding intertidal and subtidal zones (>Fig. 16.6.3). Toward the southeastern end is the Hadd ad Dib (Wolf ’s Tail), a broad spit with numerous beach ridges that narrows to an elongated ridge that passes gradually beneath the sea (>Fig. 16.6.4) and is preserved in beach rock that shows karstic weathering. Layers of thin calcareous sandstone outcrop along the relatively steep northeast flank and dip southwestward, so that the spit appears to have been tilted in that direction. On the south side, a shallow lagoon is flanked by a long shelly sand spit running southeast, then a shallow sea with a sandy floor southward in the Gulf of Salwah.

⊡⊡ Fig. 16.6.1 Dunes on the Ras al Barr at the southern point of Bahrein. (Courtesy Geostudies.)

Bahrein

⊡⊡ Fig. 16.6.2 Beach ridges on the southeast coast of Hawar Island. (Courtesy Geostudies.)

⊡⊡ Fig. 16.6.3 Sand bars off the south coast of Hawar Island. (Courtesy Geostudies.)

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⊡⊡ Fig. 16.6.4 View southeast to the Hadd ad Dib spit. (Courtesy Geostudies.)

References Doornkamp JC, Brunsden D, Jones DKC (1980) Geology, geomorphology and pedology of Bahrain. Geo Abstracts Limited, Norwich

Shinn EA (1969) Submarine lithification of Holocene carbonate sediment in the Persian Gulf. Sedimentology 12:109–144 Willis RP (1967) Geology of the Arabian Peninsula: Bahrain. U.S. Geol Sur Prof Pap 560-E. U.S. Government Printing Office, Washington, DC

16.7 Saudi Arabia, Persian Gulf Coast

Eric Bird

1.  Introduction The Kingdom of Saudi Arabia has an Persian Gulf coast about 700 km long. The Arabian Peninsula has a steep western escarpment rising to 1,500–3,000 m at the edge of a plateau that declines gradually eastward, flattening to a broad coastal plain. Longer watercourses run eastward, but there are no permanent rivers, and the landscape is largely desert. The climate is hot and arid. The Persian Gulf has summer temperatures ranging from 32°C in the south to 20°C in the north, falling to between 22 and 16°C in winter. Salinity is about 40 parts per thousand (ppt) in the south, but diminishes off the inflowing rivers in the north. Coral reefs are extensive in the southern part of the Persian Gulf, but become sparse in the cooler and more turbid northern regions. Mangroves are found in sheltered sites behind fringing reefs and beside inlets. On the Persian Gulf coast the mean spring tide range is about 1.5 m and a sandy shore is exposed at low tide. Seas in the Persian Gulf are generally slight (waves less than a metre high) and often calm, but there are occasional storms, particularly in winter, producing waves up to 4 m high. Currents are weak, and mainly wind-driven. The Late Quaternary marine transgression brought the sea to its present level about 6,000 years ago, drowning the mouths of wadis to form the long narrow and deep marine inlets called khawr on the Persian Gulf coast (Behairy 1983). These are straight or meandering, bordered by coral gardens or reefs, often interrupted at the inner end by sediment washed in from the wadi. Coral reefs that had formed in Pleistocene times and been dissected and reduced during the Last Glacial low sea level phase were recolonised by corals as the marine transgression proceeded, and the growth of Holocene reefs ensued (Gvirtzman 1977). There was a brief phase of slightly higher sea level on the Persian Gulf coast, followed by a lowering which resulted in the emergence of some coral reefs and the deposition of beaches and beach ridges behind fringing reefs. Emergence led to deposition of beaches, beach ridges and spits in front of a low cliffed coast, enclosing a number of intertidal lagoons linked to the sea through

narrow inlets. The beaches and rocky shores shelve gently to shallow sandy nearshore areas, and at low tide broad tidal flats are exposed; the sea floor is carpeted with bioclastic sandy sediment resting on an extensive sea floor rock pavement produced by Holocene submarine lithification (Shinn 1969), and there are large areas of seagrass growth. On both coasts beach rock has formed where beach sand and gravel have been cemented by aragonite. There is evidence of higher Pleistocene coastlines, includ­ing emerged coral reefs, on the Persian Gulf coasts, with dislocations resulting from uplift and faulting.

2.  Persian Gulf Coast (South to North) The Saudi Arabian coast begins at the southern end of the Gulf of Salwa, where the sea is very saline (about 70 ppt) and the shallow waters sheltered, with numerous elongated shoals and channels running south-north behind the Qatar Peninsula. The Al Jafurah desert hinterland has barchans moving southeast in response to the strong northwesterly shamal. Sand and dust are blown into the Persian Gulf (Sugden 1963). There are several large sebkhas, such as Al Uqayr on the low-lying coast that passes west of Bahrein, and curves out to the Ras Abu Urayqi Peninsula, which shelters a small gulf south of Dhahran. There are dunes, interdune hollows, sand sheets and sebkhas with quartzose sediment (Fryberger et al. 1983), as well as salt marshes and some mangroves in sheltered sites. Emerged coastlines indicate Holocene tectonic movements on the east coast of Saudi Arabia (McClure and Vita-Finzi 1982; Vita-Finzi 1991). North of the causeway that links Saudi Arabia to Bahrein there are many offshore coral reefs. The coast borders the Persian Gulf and receives waves up to 30 cm high generated by northerly winds that move beach sediment southward, forming the Ras Tannurah spit. Occasionally there are southeasterly winds generating waves that move it back to the north. There has been beach erosion and coastline retreat, but it is slow and intermittent, occurring when strong winds accompany high tide. There are large sebkhas bordered by spits and beach ridges, one opening

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northward to Al Jubayl (Johnson et al. 1978). Further north a causeway follows low narrow islands that run out to Abu Ali and shelter several bays and headlands. These culminate in the Ras az Zawr promontory, where a sandy shoal runs out into the Persian Gulf. The Manifah oil field extends offshore. Manifah borders an intricate bay which has a series of parallel ridges and narrow straits and the Ras al Tanaqib promontory to the north. Ras Saffaniya is another promontory where pipelines run out to the Saffaniya undersea oil field. Further north a series of arcuate shallow embayments and broad sebkhas are separated by cuspate forelands of slightly cemented Pleistocene calcareous oolite, which con­ tinue seaward as submerged sand spits, notably the Ras al Missaab and the Ras al Khafji. The sequence of sebkhas and forelands continues across the Kuwait border.

References Al-Sayari SS, Zotl JG (eds) (1978) The quaternary period in Saudi Arabia. Springer-Verlag, New York

Behairy AKA (1983) Marine transgressions in the west coast of Saudi Arabia (Red Sea) between mid-Pleistocene and Present. Marine Geol. 52:M25–M31 Fryberger SG, Al-Sari AM, Clisham TJ (1983) Eolian dune, interdune, sand sheet and siliciclastic sabkha sediment of an offshore prograding sand sea, Dharan area, Saudi Arabia. Bull Am Assoc Petrol Geol 67:280–312 Gvirtzman G (1977) Morphology of the Red Sea fringing reefs. Memoirs, Bureau of Geological and Mining Research, Paris, 89:480–491 Johnson DH, Kamal MR, Pierson GO, Ramsay JB (1978) Sabkhas of eastern Saudi Arabia. In: Al-Sayari SS and Zotl JG (eds) The quaternary period in Saudi Arabia. Springer-Verlag, New York, pp 84–93 McClure HA, Vita-Finzi C (1982) Holocene shorelines and tectonic movements in eastern Saudi Arabia. Tectonophysics 85:T37–T43 Shinn EA (1969) Submarine lithification of Holocene carbonate sediment in the Persian Gulf. Sedimentology 12:109–144 Sugden W (1963) Some aspects of sedimentation in the Persian Gulf. J Sediment Petrol 33:355–364 Vita-Finzi C (1991) Dating of an uplifted shoreline in eastern Saudi Arabia. Tectonomorphology 194:197–201

16.8 Kuwait and Iraq

Eric Bird

1.  Introduction At the head of the Persian Gulf the State of Kuwait has a mainland coast 225 km long, increasing to 290 km when the associated islands are included (>Fig. 16.8.1). The landscape is largely a gently undulating rocky and sandy desert, with urbanised and industrialised areas based on oil resources. The rock formations are mainly limestones and sandstones of Tertiary and Pleistocene age, the Jaz alZawr escarpment northwest of Kuwait Bay being an outcrop of Pleistocene oolitic limestone that rises 144 m above sea level, and there is a coastal fringe of Holocene beach sand and marshes. The climate is subtropical, hot and dry in summer when daytime temperatures often exceed 50°C, cool with occasional rain in winter. Kuwait has 13.5° in January and 36.6° in July, and the mean annual rainfall is 125 mm. The prevailing winds are northwesterly, strong during the shamal, but there are often southerly winds in spring and southeasterly in summer. There are occasional dust storms, particularly in winter. Tides are semidiurnal, with a mean spring tide range of 2.0 m at Ahmadi on the Persian Gulf coast, increasing to 3.5 m at the western end of Kuwait Bay and 2.6 m at the tidal mouth of the Shatt al Arab to the north. At low tide, wide areas are exposed along the shores of Kuwait Bay, sandy in the south and muddy in the north, and there are muddy intertidal areas bordering Bubiyan Island (863 sq km) and the mouth of Shatt el Arab. The muddy sediment has come mainly from the Euphrates River by way of the Shatt el Arab outflow. Waves generated by northwesterly winds generate small waves (up to 20 cm) in Kuwait Bay, while on the east-facing coast there are alternations of longshore drifting, northerly winds producing waves up to 30 cm high that move beach sediment southward and southeasterly winds generating waves that move it back to the north. There is net southward drifting of 50,000 m3/year south of Ahmadi. There has been beach erosion and coastline retreat, but it is slow and intermittent, occurring when the shamal accompanies a high tide. It is a problem locally where the coast is lined with villas built close to the beach, and rock revetments and concrete sea walls have been

constructed, but in recent decades artificial beach nourishment has been preferred to the building of solid structures. Currents are weak, and mainly wind-driven. The generally high salinity of the Persian Gulf is somewhat reduced in the northern region, in receipt of fresh water from the Shatt al Arab. The Late Quaternary marine transgression brought the sea to its present level about 6,000 years ago, and there was a brief phase of slightly higher sea level followed by emergence, which resulted in the deposition of beach­ ­es,  beach ridges and spits, enclosing a number of small lagoons and intertidal sebkhas, saline marshes linked to the sea through narrow inlets (Al-Sarawi et al. 1985; Al-Zamel and Al-Sarawi 1998). The beaches shelve gently to shallow sandy nearshore areas, and at low tide broad tidal flats are exposed (El-Sayed and Al-Bakri 1994). A thin (about 10 cm) sea floor calcrete layer, formed where sand has been cemented by carbonate precipitation, ex­­ tends as a rocky terrace offshore, and has in places been disrupted by quarrying. Beach rock has formed where beach sand and shells have been cemented by aragonite (Shinn 1969). Locally, emerged beach rock has been cut back by marine erosion to form cliffs 1–2 m high, as at Sulimiyah. Coral reefs are sparse in the northern part of the Persian Gulf, but seagrass beds are extensive. Mangroves are close to their latitudinal limits, but grow locally in inlets and sheltered areas, notably on the island of Bubiyan.

2.  The Coastline of Kuwait North from the Saudi Arabian border the coastline consists of a series of arcuate shallow embayments between cuspate promontories of slightly cemented Pleistocene calcareous oolite, which continue seaward as submerged sand spits. Ras Al-Khafji, the first cuspate promontory, has a tidal outlet from an extensive marshy sebkha, Tafal Al-Uthami, on its northern shore. An irregular beach runs along a small bay to the next foreland, Ras Bardhalaj, and on to Al Khiran, where spits enclose a small lagoon (>Fig. 16.8.2). Longshore spits have grown southward here.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_16.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 16.8.1 The coast of Kuwait. K – Kuwait City, B – Bubiyan, F – Failakka, S – Shatt el Arab. 1 Ras Al-Khafji, 2 Ras Bardhalaj, 3 Al Khiran, 4 Ras Al-Zour, 5 Ras Al Qulay’ah, 6 Ras Al-Ard, 7 Umm Al-Nemmel, 8 Jaz-al-Zawr escarpment. (Courtesy Geostudies.)

Ras Tanura Qatif

coral reef

Dammam

marsh, swamp

Manama

sand Bahrain

Halul

Hawar

Gul

Dukhan Das

w Sal f of

Doha

a

QATAR

SAUDI

Dalma

Grea t Baniyas

Pea

rl

ARABIA ⊡⊡ Fig. 16.8.2 Al Khiran lagoon. (Courtesy Geostudies.)

The sandy coast runs out to a sharp cuspate promontory, Ras Al-Zour (>Fig. 16.8.3), which carries oil tanks and a loading terminal. Offshore, Umm al Maradim, Qaruh and Kubbar are low sandy islands. A broad curving bay on the northern side runs out to another cuspate promontory, Ras Al Qulay’ah, which has dunes on its northern shore. The coast is then relatively straight northward past oil terminals and petrochemical plants at Mina Abdullah, Shua’ibah and Mina Al Ahmadi. At Ras Al-Ard the coast swings westward into Kuwait Bay, along the waterfront of Kuwait City. There is a harbour enclosure and beach compartments artificially renourished between groynes. The tide range increases westward, where a spit and outlying sand island, Umm Al-Nemmel, are the outcome of westward longshore drifting. A port has developed at Shuwaikh (>Fig. 16.8.4). The northwestern shore of Kuwait Bay is low-lying and marshy, with broad intertidal mudflats, the hinterland rising in a series of terraces that mark emerged coastlin­ ­es,  overlooked by the Jal-al-Zawr escarpment. The coast curves round to a foreland at Qaar Al-Sobiya, and then runs northward alongside a narrow tidal channel, the Khor Al Sabiyah, which separates the large low island of Bubiyan. Bubiyuan is a segment of Pleistocene delta, with plains of clay and silt and some sand ridges. The southern

Kuwait and Iraq  

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⊡⊡ Fig. 16.8.3 Ras Al-Zour cuspate foreland. (Courtesy Geostudies.)

⊡⊡ Fig. 16.8.4 Shore east of Kuwait City, showing low cliff of emerged beach rock and a narrow beach descending to a gravelly foreshore (dissected calcrete). (Courtesy Geostudies.)

coast has been cut back by wave erosion, and sandy material has drifted alongshore to form a narrow spit that shelters a prograded marshy eastern coast. The northern part of the island has many intersecting tidal creeks, the outcome of subsidence, and alongside these are stunted

mangroves. At low tide mudflats are exposed as the sea subsides into convergent deeper channels. Failakka, to the south, is a low-lying limestone island (12 km by 6 km) with sandy beaches, and is much used by Kuwaitis for recreation.

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3.  The Coastline of Iraq A short sector (58 km) of the northern coast of the Persian Gulf lies within Iraq. The area is dominated by the large swampy delta of the Euphrates, Tigris and Karun Rivers, from which the Shatt el Arab is the main outflow channel. Deposition of sand, silt and clay from these rivers produced deltaic features during Pleistocene high inter­ glacial  sea level stages, and these have subsided, and been  partly buried by similar Holocene delta sedimen­ tation. The shores are backed by salt marshes grading into  extensive freshwater marshes, and there is a wide accreting intertidal zone, particularly off the mouth of the  Shatt el Arab. Near the river mouth a pier leads out  across the intertidal mudflats to an oil terminal at

Khor-al-Amaya. The border with Iran runs alongside the Shatt el Arab channel.

References Al-Sarawi M, Gundlach ER, Baca BJ (1985) Kuwait, an atlas of shoreline types and resources. Foundation for the Advancement of Science, Kuwait Al-Zamel A, Al-Sarawi M (1998) Late Quaternary sabkha sedimentation along Kadmah Bay coast, Kuwait, Arabian Gulf. Arab Gulf J Sci Res 16:471–495 El-Sayed MT, Al-Bakri D (1994) Geomorphology and sedimentary/ biosedimentary structures of the intertidal environment along the coast of Kuwait, north-west Arabian Gulf. Geologische Rundschau 83:448–463 Shinn EA (1969) Submarine lithification of Holocene carbonate sediments in the Persian Gulf. Sedimentology 12:109–144

16.9  Iran

Rodman E. Snead

1.  Introduction The Iranian coastline is about 2,440 km long. A rugged escarpment at the edge of an interior plateau overlooks a coastal area of Tertiary rocks with ranges produced by Alpine folding and intermittent narrow coastal plains. The coastline runs parallel to the anticlinal ridges, one of which runs through the island of Qeshm. In the south (Laristan), there are numerous circular salt domes formed by diapiric intrusions from an underlying Cambrian salt horizon. Rivers that drain to the coast have short, steep courses. The landscape is largely desert. The climate is hot and arid in summer and mild and rainy in winter. Mean annual rainfall decreases along the coast from 277 mm at Bushi­­re  to 147 mm at Bandar Abbas and less than 130 mm at Jask and Chah Bahar. Spring thunderstorms cause sudden flooding in rivers draining to the coast, and floods leave thick deposits of mud on alluvial flats around river mouths. The Persian Gulf has summer temperatures ranging from 20° in the north to 32° in the south, falling to between 16° and 22°C in winter. Salinity is low off the Shat-el-Arab, which discharges water from the Tigris and  Euphrates Rivers, but increases to about 40 parts per thousand in the south, where high evaporation results in an inflow of Indian Ocean water through the Strait of Hormuz. Coral reefs occur at intervals along the coast, becoming extensive in the southern part of the Persian Gulf. Mangroves are found in sheltered sites behind fringing reefs and beside inlets. There are oilfields north of Bushire and beneath the Persian Gulf to the south. Tides on the Persian Gulf coast of Iran (for the Caspian coast of Iran >  Iran-Caspian Sea Coast) are generally small. At the Shatt el Arab outer bar in the north and at Khowr-e Musa mean spring tide range is 2.6 m. Bushire has 1.5 m, Bandar e Lengeh 2.1 m, Bandar Abbas 2.6 m, while on the south-facing Gulf of Oman coast Chah Bahar has 1.8 m. Seas are generally slight (waves less than a metre  high) and often calm, but there are occasional storms, particularly in winter, producing waves up to 4 m high from the west or northwest. Wave action becomes

stronger on the Makran coast of Iran, where southerly swell from the Indian Ocean arrives through the Gulf of Oman. Currents are weak, mainly wind-driven, in the Persian Gulf. The Late Quaternary marine transgression brought the sea to its present level about 6,000 years ago, drowning the mouths of valleys to form the marine inlets. There was a brief phase of slightly higher Holocene sea level, followed by a lowering, which resulted in emergence and the deposition of beaches, beach ridges, and spits in front of a low cliffed coast, with a number of intertidal sebkhas linked to the sea through narrow inlets. The Persian Gulf coast beaches shelve gently to shallow sandy nearshore areas, and at low tide sandy flats are exposed. The sea floor is carpeted with bioclastic sandy sediment and there are large areas of seagrass growth. Beach rock has formed where beach sand and gravel have been cemented by aragonite. There is evidence of higher Pleistocene coastlines, including emerged coral reefs, with dislocations resulting from uplift and faulting. Higher Quaternary coastlines have been studied in southern Iran, particularly on the Makran coast (Vita-Finzi 1981; Reyes et al. 1998).

2.  The Iranian Coastline From the Iraq border along the Shatt el Arab, the outflow channel from the Tigris and Euphrates Rivers, the coastline runs round the eastern part of a swampy delta to Khowr-e Musa, an intricate embayment in an area of subsidence at the mouth of the Jarahi River. To the east is the Zohreh (Zuhr-el-rud) delta, fringed by extensive mudflats where tide range is up to 4 m. The coast then swings southward past Bandar-e Deylam to Bandar Rig, where a spit had grown southward. To the southwest is the Island of Kharg, bordered by cliffs and rising to 83 m above sea level. A smaller island, Kharkhu, lies to the north. On the mainland coast is the large lobate Helleh delta, with extensive swamps that extend behind the bay to the south, which is an area of rapid sedimentation with mudflats and

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mangrove-fringed marshes. Bushehr (Bushire) is a port town standing on a triangular peninsula, part of a promontory that extends southeast to Ras e Halileh. A coast road runs south to Rishahr, which has the ruins of an ancient city. A narrow sandy plain fronts a ridge parallel to the coast, which is interrupted by the valley of the Mand, where the meandering Mand River has built a large delta of irregular outline. Longshore drifting is southward, and has formed spits along the coast. At Kangan, a fishing village, the coastal plain narrows, dominated by high scarped mountains, except where bays correspond with subsidence basins or synclinal areas, as in Naband Bay. A generally steep coast continues southeast to Bandar-e Maqam. A chain of islands offshore include Shuaib, which rises to 37 m, Kish, a beach-fringed island rising to 42 m and the little islet of Tumb. Bandar-e Taheri has the ruins of an eighteenth century fort. Hendroabi and Geys are outlying high islands, and the coast continues to Bander-e Lengeh, where it swings ENE. The elongated island of Qeshm has anticlinal ranges and rocky coasts. It is separated from the mainland by Clarence Strait, which has tidal mudflats with some mangroves. Bandar Abbas, a long narrow urbanized strip, stands behind a bay to the east, and is backed by hilly country in which tectonic deformation continued into Holocene times (Vita-Finzi 1979). Beyondit, the deltaic plain of the Baghu and Shirin Rivers has muddy shores,

and long discontinuous spits parallel to coastline enclose lagoons. Offshore is the high island of Hormuz, developed on a salt dome, rising to White Peak (235 m), with steep southern coast and a northern lowland running out to a point on which a ruined Portuguese fort stands. In the northwest is the site of an ancient city. The Strait of Hormuz curves round the Musandam Peninsula of Oman, and the Iranian coast swings southward past the Minab River delta to Sirik and Kuh-Mabarak and the Ras al Kul headland, where the coast turns eastward along the Gulf of Oman. In the hinterland are the Makran Ranges, where Tertiary sandstones, marls, and clays have been strongly folded and faulted in a series of anticlines and synclines, uplifted as sharp-crested ridges, scarps, and platforms, largely bare of vegetation and dissected by gullies (Shearman 1976). The coastal plain is ­seldom more than 32 km wide, and includes uplifted and  tilted marine terraces (Reyss et al. 1998). It is gen­ erally  arid, with locally a sparse growth of bunch grass and woody shrubs, but after winter rains there is a brief appearance of grasses and herbs on the usually dusty plains. Jask stands out on a long sandy promontory, with a low arid terrace cliffed along the coast to the east (>Fig.  16.9.1). The sandy coastal plain is interrupted by protruding scarps and dip-slopes. The Jagin River reaches the coast through a series of channels converging in a coastal lagoon behind a barrier of bare, drifting dunes.

⊡⊡Fig. 16.9.1 Cliff cut into low terrace near Jask.

Iran

⊡⊡ Fig. 16.9.2 Cliffed terraces at Chah Bahar.

⊡⊡ Fig. 16.9.3 Cliff cut in tilted terrace east of Chah Bahar.

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⊡⊡Fig. 16.9.4 Cliffed coast and capping dunes east of Chah Bahar.

⊡⊡Fig. 16.9.5 Cliffed coastal terrace backed by mountain scarp west of Briz.

The coastal plain to the east is crossed by the Gabrik and Sadij Rivers. A barrier spit has grown from west to east to enclose a lagoon near Poshti, and the coast then becomes a series of headlands and bays with increasing exposure to southerly swell from the Indian Ocean. This is the barren Makran coast of Iran (Snead 1970), where coastal terraces show evidence of Late Quaternary uplift and tectonic deformation (Vita-Finzi 1981) and earthquakes occur. It is well shown on Google Earth. A rocky promontory at Ras Tank is attached to the mainland by a tombolo with bare dunes, and farther east the Ras e Puzm promontory has a western bay backed by mobile dunes. This promontory borders Chah Bahar Gulf, semi-­ circular with bay-head beaches receiving refracted swell and an eastern headland that is cliffed at the edge of a broad terrace (>Fig. 16.9.2). To the east cliffs have been cut into tilted terraces incised by wadis (>Fig.  16.9.3), with cliff-top dune ridges oriented at right angles to the coastline (>Fig. 16.9.4). There are some short rivers with mouths blocked by a long sandy beach, and narrow coastal terraces backed by steep scarps and trimmed seaward by receding cliffs (>Fig. 16.9.5). The irregular coastline continues to Ras-e-Fasteh on the western side of Gavater Gulf, where the terrace has been tilted landward and low cliffs border the coastal terrace toward Gavater (>Fig. 16.9.6). North of Gavater, the Dashtiari River has built a delta into a narrow inlet, and there are patches of mangroves on intertidal mudflats. The Pakistan border crosses the coastline at the head of the bay.

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⊡⊡Fig. 16.9.6 Cliffed low terrace at Gavater.

References Reyes JL, Pirazzoli PA, Haghipour A, Fontugne M (1998) Quaternary marine terraces and tectonic uplift rates on the south coast of Iran. Geol Soc Spec Pub 146:225–237 Shearman DJ (1976) The geological evolution of southern Iran. Geogr J 142:393–410

Snead RE (1970) Physical geography of the Makran coastal plain of Iran. Geography programs, Office of Naval Research, Washington, DC Vita-Finzi C (1979) Rates of Holocene folding in the coastal Zargos near Bandar Abbas, Iran. Nature, 278:632–634 Vita-Finzi C (1981) Late Quaternary deformation on the Makran coast of Iran. Z Geomorphol, Suppl(Bd.) 40:213–226

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Mountains produced by Alpine earth movements extend eastward, then northward in Pakistan, and eastward as the Himalayas before turning southward to Burma. The Indian subcontinent juts southward from this mountainous zone, between the Arabian Sea and the Bay of Bengal. The coastline from Pakistan to Burma is subject to swell from the Indian Ocean and the waves generated by monsoon winds, southerly and south-westerly in June to September and north-easterly in October and November. The > Pakistan coast has sandy beaches bordering a narrow coastal plain backed and interrupted by marine terraces on Tertiary sandstones and limestones, which form cliffed headlands where they reach the coast. The Makran, Las Bela and Karachi coastal regions are neotectonic, subject to recurrent earthquakes. To the east is the large Indus delta, with tidal estuaries between beaches, spits and barrier islands, extensive mudflats and some mangroves (Nayak 2005). East of the Indus delta the Rann of Kutch is a saline desert, with an indented coastline of tidal creeks and mudflats. In > India the Kathiawar coast is also low-lying, but east of the Gulf of Khambhat the coastal plain is backed by the escarpment of the Western Ghats, which extends behind the whole of the west coast of India. South of Mumbai the coastal plain is generally about 40 km wide. There are sandy beaches separated by numerous rocky promontories (basalt in Maharashtra and granite-gneiss further south) and interrupted by estuaries and tidal inlets. In Goa the beaches are longer, backed by sand dunes and estuarine lagoons, and similar features are seen further south in Karnataka. Sectors of cliffed coast are backed by emerged terraces, and fronted by stacks and islands. Longshore drifting is northward, particularly during the boisterous SW monsoon. Mean spring tide range is up to 11 m in the NW but diminishes to 3.6 m at Mumbai and about 1 m south of Mangalore. The east coast of India has a wider coastal plain, with several large deltas. There are occasional rock outcrops on cliffed hilly sectors with caves and shore platforms, but much of the coast is fringed by sandy beaches that become barriers in front of lagoons such as Pulicat Lake and Chilka Lake. Spits have grown northward. There are many

mangrove-fringed tidal creeks, particularly at the mouths of distributary channels on the Krishna and Godavari deltas. Further north, the large Ganges delta has numerous mangrove-fringed tidal creeks and islands, including the extensive mangrove forests in the Sunderbans. > Sri Lanka has a core of Pre-Cambrian rock fringed by Miocene limestones and Quaternary deposits. Much of the coastline consists of sandy beaches, interrupted by rocky headlands. Beaches are generally backed by dunes, and there are numerous lagoons, particularly on the east coast. Coral reefs are extensive. In addition to Indian Ocean swell, wave action is related to the SW and NE monsoon winds. The SW, S and E coasts are exposed to  strong wave action, but the NW is more sheltered because of the shallowness of the Gulf of Mannar and Palk Strait, and has spits and barrier beaches, estuaries, tidal mudflats and mangroves. Mean spring tide range is less than 1 m. Much of the >Bangladesh coast is low-lying and swampy, including the mangroves of the Sundarbans and the inlets and islands of the eastern part of the Ganges delta. The tide range is between 2.5 and 4.5 m, and there is tidal scour in the inlets and channels. Wave action is generally weak because of the wide shallows at the head of the Bay of Bengal, but there are occasional storm surges accompanying tropical cyclones that cause extensive flood­ ­ing and erosion. The Chittagong coast to the east has some cliffed sectors between beaches fronted by tidal sandflats, with many shoals and islands. In >Myanmar (Burma) (the mountains formed on Mesozoic and Tertiary rocks folded and faulted by Alpine earth movements run N–S and influence the outlines of the coast. Long ridges with steep coasts run out between narrow straits at the mouths of river valleys, and there are many high islands. Mean spring tide ranges are from 2 to 6  m. Wave action becomes stronger southward as the nearshore area deepens, and there are cliffs on protruding headlands. Mangroves fringe estuaries and sheltered inlets, and there are patchy coral reefs. Beaches occupy bays, and mud volcanoes form temporary islands, which are soon reduced to shoals by wave action. To the south, the Irrawaddy delta has an intricate E–W coastline with many

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_17.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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mangrove-fringed channels and tidal creeks: the tide range is about 5 m. From the Gulf of Martaban the Tenasserim coast runs southward and the mountains decline to many elongated peninsulas with steep slopes and cliffs, breaking into chains of high islands. Interven­ ing straits and inlets have mangroves and wide inter­ tidal  mudflats. There are bay beaches, some backed by dunes. Ocean swell and waves generated by the SW ­monsoon break heavily on the outer shores of islands

and ­peninsulas, but the intricate configuration shelters much of the coastline.

Reference Nayak GN (2005) Indian Ocean coasts, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 554–557

17.1 Pakistan

Rodman E. Snead

1.  Introduction The coastline of Pakistan is about 1,365 km long from the Iranian border eastward to the Indian border at the Rann of Cutch. Much of the coast has been influenced by the uplift of marine platforms and the Makran mountain ranges. The Las Bela Valley represents a structural trough forming part of the Owens Fracture Zone that extends southward under the Arabian Sea. Both the southeastern part of the Indus delta and the Rann of Cutch appear to be undergoing subsidence. The Pakistan coast shows evidence of sea level changes (Ali Khan et al. 2002) and is one of the most tectonically active regions in the world, with 20 recorded earthquakes between 1939 and 1960. The November 1945 earthquake was particularly severe: it generated a tsunami, reported to have been 12–15 m high, across the coast, and a section of the shore near Pasni was uplifted by 4.6 m. This coastal region comprises long stretches of sandy beach backed by wide valleys or largely barren coastal plains. The low-lying areas are interrupted in places by uplifted terraces that form hammerhead peninsulas, backed by hills and mountains. The Makran ranges, with elevations of 914–1219 m, rise abruptly some 32 km inland from the coast. A narrow continental shelf borders most of the Pakistan coast except off the Indus delta. Scattered mudbanks and shoals are found along the coast, most of them within 1.6 km of the shore. A concentration of islets, rocks and shoals clusters around Astola Island some 40 km southeast of Pasni. The coastal climate is semiarid, with monsoonal rain increasing in the east. There is a dry season between March and May. Karachi has mean monthly temperatures of 16.1°C in January and 30°C in July, with an average annual rainfall of 196 mm. Most of the Pakistan coast,  especially west of Ras Muari (Cape Monze), is exposed to southwesterly ocean swells that become very choppy ­during the southwest monsoon months, June through September. An occasional Arabian Sea cyclone (a small but intense hurricane) creates severe erosion and flooding along the coast. Mean spring tide ranges on the ocean coastline are ­generally 0.9–1.8 m with slightly

higher ranges (up to 2.7 m) around Ras Muari: Pasni has 1.8 m and Karachi 2.3 m. Beach sediment are predominantly sandy, and longshore drifting is generally from west to east, but in the scallop-shaped bays the refracted ocean swells direct sand towards the centre, forming long barrier spits and bars. Beach accretion in the shelter of former islands has formed the large tombolos that link the Gwadar and Ormara headlands with the inner coastal plains. West of Manora Point the beach sands are mainly quartzose, while those east of Karachi harbour and fringing the Indus delta are highly micaceous. Coastal dunes are frequent mainly on the east side of the Las Bela Valley, east of Karachi harbour at Clifton, and along large sections of the low, very dry, windswept Makran coastal plain. More stabilised Pleistocene sand deposits are extensive on the east side of the Las Bela Valley. Most of the rivers are small and intermittent, sinking into the sands of the coastal plains before they reach the sea. The three larger rivers, the Hab, the Porali, and the Dasht, have very little discharge, except during floods, and drain into estuarine lagoons or shoal-filled river mouths. Except for Karachi harbour, which is dredged, there are no good ports along the entire Pakistan coast. In fact, the coast is largely uninhabited, except for eight or ten small fishing villages. Ancient seaports, however, have been under investigation by archaeologists. Several of these 3,000-year-old sites lie as far as 32 km inland from the present coast. The Pakistan coast may be divided into four sections from west to east: the Makran coastal region, the Las Bela coastal plain, the Karachi region and the Indus region.

2.  The Makran Coast In Gwatar Bay, near the Iranian border, is a low swampy region that forms the delta of the Dasht River. The coastal plain is 48 km long and extends inland for more than 241 km. The Dasht brings little sediment to the coast. In this arid region (average annual rainfall: 127 mm), it is an intermittent stream that floods only once every 2 or 3 years.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_17.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Most of the water sinks into the sands and gravels of the river bed or is trapped in the shallow lagoons behind the coastal barrier bars and spits. The Makran coast, between the Jiwani headland and Ras Malan, consists of long stretches of sandy shore backed by valleys or wide coastal plains with numerous beach ridges, notably to the west of Pasni. There is a bold headland at Gwadar (>Fig. 17.1.1), linked to the mainland by a broad sandy beach and another tombolo attaching the headland at Ormara. Between Pasni and Ormara the sandy beaches form a barrier separating the Kalmat Khor Lagoon from the sea. These low-lying areas are interrupted in a  number of places by uplifted and tilted terraces (>Fig. 17.1.2) that form hammerhead peninsulas backed by hills and mountains. This section makes up the bulk of the Pakistan coastline, about 473 km. Inland are segments of the Makran Ranges, which in general lie about 32 km from the coast except at Ras Malan, where the Hinglaj Range comes to the shore as a massive sandstone and limestone headland on a coast up to 600 m high. West of the low Mor and Haro ranges, the Makran hills turn southwest and at Ras Malan begin to trend east–west parallel to the coast. The Ham, Mor and Kirthar ranges to the east trend north–south and come to the sea as headlands at Ras Muari and Gadani. The Makran ranges are composed mainly of Oligocene flysch, poorly indurated sandstones, mudstones and shales, tightly folded into anticlinal and synclinal ridges. Severe erosion results in a bizarre landscape with large cap-crowned

pillars and pedestals that rise abruptly above jagged ridges and ravine-ridden lower hills of clay. The spectacular marine platforms of Jiwani, Gwadar and Ormara are quite different from the inner hills. On the seaward side they are made up predominantly of PlioPleistocene marine conglomerates and shelly limestones, while the inner, landward-facing side consists of Miocene– Pleistocene mudstones and sandstones. These platforms represent large fault-blocks, or horsts, that are tilted seaward with a dip of 3–5°. Ormara, which is higher than Gwadar and Jiwani, has elevations of 183–305 m on the seaward side and 427 m on the inner side. The Gwadar and Ormara platforms, formerly islands off the coast, are now connected to the mainland by large sandy tombolos. The Ormara tombolo is 3.2 km wide and 10 km long. From the top its growth, through a series of accretion ridges, can be clearly seen. Between the rocky headlands control points are found on scallop-shaped bays or long barrier islands with shallow lagoons. One of the largest of these lagoons is Kalmat Khor, to the west of Ormara (Snead 1968).

3.  The Las Bela Coastal Plain This section of coast makes up the eastern part of the Makran coast of Baluchistan. It extends from Ras Malan eastward 260 km to Ras Muari (Cape Monze), and inland to include the southern part of the triangular Las Bela plain

⊡⊡ Fig. 17.1.1 The headland at Gwadar, showing salt mounds and pans.

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⊡⊡ Fig. 17.1.2 Tilted terraces cut by cliffs at Ras Malan.

and the foothill regions of the low Mor and Haro mountain ranges. These two ranges, about 97 km apart at the coast, merge 113 km to the north to enclose the valley. Ras Malan bounds the Las Bela coastal plain on the west. It is a spectacular tableland of sandstone and shelly limestone cut into irregular segments by 300 m gorges. The highest point on this tableland is 610 m above sea level, and sheer cliffs, 457–600 m high, drop directly to the sea. The fault scarp itself is 4.8 km long. On the west side of the Las Bela valley, near the southern end of the Haro Range, is a group of mud volcanoes. The largest, called Chandragup, is an almost perfect cone less than two miles from the Arabian Sea; it reaches a height of 58 m. About a dozen small, white mud volcanoes, of which the highest is some 32 m, lie 3.2 km east of Chandragup, and a single cone, 20 m high, lies about the same distance to the west. These mud volcanoes represent rapid sedimentation and widespread strike-slip faulting of Tertiary mudstones, sandstones and shales. The cones occur where plastic clay, overlain by heavier silts and sands, reaches the surface through vents and flows out as viscous mud. The main coastal portion of the Las Bela plain comprises a series of beach ridges, shifting sand dunes (>Fig. 17.1.3), two long barrier bars with hooks at their ends, and the tidal flats around Miani Lagoon. The largest mangroves along the Pakistan coast grow in this lagoon. Both Rhizophora conjugata and Avicennia alba can be found, with individual trees reaching up to 9 m.

Several headlands mark the eastern side of the Las Bela valley. The Kirthar ranges reach the coast at Ras Muari, where the headland is made up of massive limestones Cretaceous to Oligocene in age, with some intrusions of Deccan lava. Marine terraces occur near Ras Muari, of which the highest attains an elevation of about 76 m and two lower ones reach heights of 15–30 m. A small Cretaceous limestone headland at Gadani reaches a height of 87 m (>Fig. 17.1.4). Wave erosion at the base of this headland has cut sea caves, blowholes and picturesque rocky promontories. Between Ras Muari and Gadani is the mouth of the Hab River. This intermittent stream is confined by a cliffed coast on the south and by Pleistocene sand deposits on the north. Although the Hab is one of the longest rivers in eastern Baluchistan, with a considerable discharge during floods, it does not have a delta. At its mouth, small compound bars have formed wide, sandy beaches. Shallow lagoons are found behind the bars.

4.  The Karachi Region The 48-km-long coastline in the Karachi region, from Ras Muari (Cape Monze) to Clifton Beach, consists of about 24 km of sandy barriers and beaches bordering the harbour of Karachi and about 24 km of low, rocky cliffs at the southern end of the Kirthar Hills. The cliffs, which are 12–15 m high on the flanks of the Kirthar Hills, become

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⊡⊡ Fig. 17.1.3 Migrating sand dunes near Siranda Lagoon.

⊡⊡ Fig. 17.1.4 Headland at Gadani.

lower to the east, away from the mountain front. They represent the eroded seaward side of a long pediment made up of Tertiary sandstones, conglomerates and shales covered by a thin veneer of gravels and sands. Coastal terraces have been tilted seaward east of Ras Muari (>Fig. 17.1.5). There are several small rocky capes with shallow scallop-shaped bays and a number of sea caves (>Fig. 17.1.6), stacks and arch. This is a complex coastal region, embracing the deeply dissected pediment

surface, uplifted marine terraces and recent oyster beds that lie 2.4–3 m above the present sea level. A 15-km-long sand spit connects the rocky headland of Manora Point with the mainland. In places this sand barrier is less than 300 m wide, but it is nevertheless the main resort beach of Karachi. Behind it is a shallow tidal lagoon 8 km long and 4 km wide. Human activity has greatly changed the tidal flats in this lagoon by building a series of salt evaporation ponds and by filling in sections

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⊡⊡ Fig. 17.1.5 Sloping terrace east of Cape Monze.

⊡⊡ Fig. 17.1.6 Caves cut in sandstone cliffs at Manora.

for urban–rural expansion. Karachi harbour has been made by dredging, widening and improving the entrance and lower part of the small, intermittent Chinna Creek. Manora Point, a perpendicular cliff, is the southeastern extremity of a narrow hill, 26 m high, which forms the southwestern side of the entrance to the harbour. The lower harbour is nearly 0.8 km wide, but the navigable

channel is reduced to 274 m by the banks that extend from either side. The upper harbour stretches northward and northeastward from the southern end of East Wharf for about 3.2 km and has a navigable width of 274–366 m for most of its length. On the incoming tides during the southwest monsoon a considerable swell rolls into the outer harbour.

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⊡⊡ Fig. 17.1.7 Low mangroves (Avicennia alba) on the Indus delta.

5.  The Indus Delta Region The Indus delta encompasses about 200 km of the Pakistan coastline. The landforms are uniform along this entire section. The region consists largely of tidal channels, most of which are remnant courses of the Indus River. Between these channels are extensive mudflats consisting of fine delta silts and clays with a high percentage of micaceous sand. Where the delta meets the Arabian Sea, the fine materials are carried offshore and coarser sands are left to form a series of barrier bars with spits and hooks at the ends. The beaches along the outer bars are some 60–120 m wide, with a gradient of 1–3°. The individual bars number about 15, and they are separated by tidal channels that vary in width from a few hundred metres to more than 1.5 km. Most of the sand bars are barren and flat, but in places low dunes reach a height of 3–5 m. The sand bars and delta channels are continually changing as tidal currents, waves and channel floods move the fine white sands along the beaches. To the south, where the Indus River is now active, mudflats extend to the coast, and there are few barrier bars. The delta merges southward into the mud and salt wastes of the Rann of Cutch (where there is a disputed border with India). At high tide the coastal region is submerged for as much as 4.8–6.4 km inland because the land is so flat. During the monsoon season, when high tides and Indus River floods coincide, the delta is submerged up to 32 km from the coast. Tides ascend the Indus River for 97 km to

Tatta, and at times form a tidal bore that is destructive to small boats. When the tide recedes and the old creeks dry, salt covers the surface or is buried under mud. There have been many changes in the configuration of the Indus delta. The seven mouths enumerated by the Roman historian Ptolemy have altered through the centuries as former channels have been choked and new ones opened up. The region is furrowed by these ancient channels, some still intact, others more or less obliterated. Because the Indus delta coastline is under constant and heavy wave attack, the currents continually redistribute sediment, mainly southeast towards the Rann of Cutch, where winds pick up the fine material and carry it inland in the form of large dunes. The Rann of Cutch itself, a rift zone uplifted on the north side and subsiding on the south, is a highly unstable region. The 1819 earthquake resulted in the submergence of about 500 km2 of low-lying deltaic land. Although a few small mangrove shrubs, mainly Avicennia alba, can be found along the tidal channels, severe flooding and deposition of delta silts limit their growth (>Fig. 17.1.7).

References Ali Khan TM, Kabir A, Sarker MA (2002) Sea level variations and geomorphological changes in the coastal belt of Pakistan. Marine Geodesy 25:159–174 Snead RE (1968) Physical geography reconnaissance: West Pakistan coastal zone. University of New Mexico Press, Albuquerque

17.2 India

G. N. Nayak · P. T. Hanamgond

1.  Introduction India has a coastline that is more than 7,500 km long, including islands (Nayak 2005). The existing coastal geomorphology has evolved largely during and since the postglacial marine transgression (Baba and Thomas 1999). There have been sea level fluctuations during the past 6,000 years, with a marked regression between 5,000 and 3,000 years bp. Storm surges, cyclones and rare tsunamis have modified coastal landforms, as have the tropical monsoon climate with its seasonal variations, and wave and current regimes. There are extensive sandy beaches, some straight, others curved, often backed by beach ridges and dunes, especially along the central west coast. In the nearshore zone there are bars and shoals. Longshore drifting has nourished spits, and barriers have formed in front of lagoons. Rivers have built deltas, and there are intertidal flats, salt marshes and mudflats, with and without mangroves. There are cliffs with and without shore platforms. The ­pioneering study of Ahmad (1972) has been followed by numerous research publications dealing particularly with coastline changes, notably those resulting from the 2004 Indian Ocean tsunami, which had a strong impact on the east coast. The Space Application Centre (SAC 1992) has carried out a comprehensive study of the Indian coast using LANDSAT and IRS data. The Indian coast has a tropical monsoon climate. Mumbai (Bombay) has a mean annual rainfall of 1,839 mm with a mid-year wet season, and Chennai (Madras) 1,257 mm with a maximum in October and November. Both have a mean minimum temperature of 24.5°C in January. Much of the coast is exposed to ocean swell arriving from the south across the Indian Ocean and refracted to move towards the west coast from the SW and the east coast from the SE. In addition, waves are generated by onshore winds, notably the NE monsoon on the east coast and the SW monsoon on the west coast. Mean spring tide range is large in the northwest, 8.8 m at Bhavnagar and 3.6 m at Mumbai, but much of the coastline is micro-tidal (Fig. 17.2.1).

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⊡⊡ Fig. 17.2.1 Holocene beach rock exposed on the Gujarat coast.

Much of the NW coast, from Bharuch to Mumbai and Ratnagiri, is backed by Deccan volcanic traps. The basaltic rocks have columnar joints that guided the formation of numerous caves and bays on the coast between Daman and Mumbai, and have disintegrated into lava boulders. The coasts along Maharashtra, Goa and north Kar­ nataka have pocket beaches flanked by rocky cliffs, estuaries, bays and mangroves in some places. Shore platforms occur on Deccan basalts along the Maharashtra coast and there are long beaches in southern Goa and north Karnataka. There are some islands predominantly of granitic gneiss (off Karwar), quartzite (off Sindhudurg) and basalt (off Mumbai). Beach dynamics are related to variations in monsoon climate along the Karwar and Vengurla coasts (>Fig. 17.2.2) (Hanamgond 2007). Mumbai stands on coastal lowland beside a broad mangrove-fringed estuary, where spurs of mangroves run out onto tidal mudflats. Evidence of recent coastal subsidence of 6–12 m in the Mumbai region was found during excavations at Princess Dock, which showed trunks of trees (some of which never grow below highest tidal level) in upright position with roots at a depth of 6 m below sea level, and at Alexandra Dock tree stumps were found 12 m below highest tidal level. The Western Ghats (in Maharashtra), the great scarp overlooking the western coast, are built of Deccan lava, which erupted in the Eocene. The coastline runs roughly parallel to the Western Ghats, and a plain of marine denudation, now bearing low-level laterite extending to the foot of the scarp, suggests that the lava plateau formerly extended much farther

west beneath the Arabian Sea. Later faulting led to the ­subsidence of the western segment, allowing the sea to reach the scarp foot. There is biological evidence (mollusca on the Western Ghats allied to the littoral marine genus Littorina) to suggest that the Ghats were once washed by the Arabian Sea. A great cliff was a feature of the  Tertiary coast, probably contemporaneous with the Narmada and Tapti rift valleys. After the faulting produced the Western Ghats there has been more recent uplift to form the low-lying Konkan strip. The Maharashtra coast, between Mumbai and Goa, shows evidence of Late Quaternary submergence followed by Holocene emergence. Radiocarbon dating of one of the emerged areas gave a date of 6460 bp. Erosion is occurring at points along this coast, particularly during the SW monsoon. The fishing village of Devbagh, 8 km south of Malvan, is one of many threatened by marine erosion, countered by building sea walls. Shore platforms on Harihareshwar coast, caves on the Guhagar coast, cliffs and rocky coasts in Kelsi and Dabhol have been studied with respect to sea level fluctuations by Karlekar (1993) (>Fig. 17.2.3). Pre-Cambrian crystalline gneiss, schist and granite occur along the Goa and north Karnataka coasts. Generally they lie behind a coastal plain ribbon of alluvium or Tertiary deposits, but locally they extend to the coast to form cliffs and rocky shores. Submergence has produced an indented coastline, with many small inlets and rias. Mudflats are extensive, within estuaries along Goa and northern Karnataka. The southern Karnataka coast has long straight beaches, and estuaries with islands

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⊡⊡ Fig. 17.2.2 Broad beach at Kille Nivati on the Maharashtra coast.

⊡⊡ Fig. 17.2.3 Tors in Metamorphic rock on the Goa coast.

and mangroves. Most of the estuaries have spits growing northward (SAC 1992) (>Fig. 17.2.4). The coastline from Bhatkal, south of Uttara Kannada, becomes smoother in outline, and has beach-fringed

s­ ectors extending past Mangalore and Calicut. The Kerala coast has marine cliffs cut into terraces 3–5 m above sea level. This coast was receding at a rate of up to 6 m/year, and comparison of 1850 and 1966 maps shows that much

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⊡⊡ Fig. 17.2.4 Steep coast with rocky shores on the Karnataka coast.

of the Kerala coast receded during that period (Earat­ tupuzha and George 1980). Walls for coast protection were constructed along 64 km of eroding coast where the sea threatened to destroy railway lines, national highways, canals and paddy fields, and narrow barriers fronting coastal lagoons. The beaches of southern Kerala are backed by dunes, which have been trimmed back by coastal erosion locally, especially near Trivandrum. There are estuarine backwaters, an ecosystem that supports fisheries and transient seasonal mud banks offshore, which may act as barriers to coastal erosion along this coast. Offshore are the Lakshadweep (formerly Laccadive) and Minicoy islands, which consist of coral reefs and lagoons, with some coral reefs rising to 2–3 m above the present sea level.

3.  East Coast of India The coast curves SE to Cape Comorin, where the PreCambrian rocks of the hinterland run through to the coast to form cliffs and rocky shores. Quaternary sediment in the Cauvery region of Tamil Nadu and in the Kolanka– Masulipatnam region of Andhra Pradesh indicate a higher sea level during the early part of the Pleistocene. Coral reefs on the Rameshwaram peninsula appear to have formed around 6000 bp. The SE coast of India is relatively sheltered from strong wave action by Sri Lanka. Several small deltas form the coastline of the Gulf of Mannar, and a long narrow peninsula, breached by Pamban Channel, runs out to

Rameswaram Island and the submerged Adam’s Bridge to Sri Lanka. The coast of Palk Strait also has low wave energy, with several small deltas such as that of the Vellar. The coast curves out to Point Calimere. To the north is a long straight coastline where strong waves generated by the NE monsoon have truncated the Cauvery delta. The seaport city of Madras stands on a broad sandy coastal plain, formed by Holocene progradation of the Coromandel coast. Numerous parallel beach ridges with some low dunes mark stages in this progradation. There are beach ridges on the Krishna delta, 4–6 m above mean sea level and 20 km inland. Each of the east coast deltas shows progradation (>Fig. 17.2.5). East-flowing rivers carry large quantities of water and sediment to the Bay of Bengal, and have built large deltas at their mouths. The NE monsoon is weak, but there are periodic cyclones and associated floods. On the Godavari delta, Sambasiva Rao and Vaidyanad­ han (1979) identified four strandlines at successively lower levels from 8 to 2 m, believed to be of Holocene age. Godavari Point, a spit enclosing Kakinada Bay, has been prolonged northward by longshore drifting by about 12 km in 100 years. The growth of the spit is mainly due to in­­ creased fluvial sediment discharge owing to deforestation. Near Vishakhapatnam there are cliffs of Pre-Cambrian rock. The Krishna delta is growing by the addition of spits and bars (SAC 1992). Chilka Lake is a large coastal lagoon partly enclosed by  a long sandy barrier island (Sriharikota Island), with marine entrances at the southern (Pulicat) and northern (Armagon) ends. Pulicat Lake is brackish, and has­

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⊡⊡ Fig. 17.2.5 Longshore bars on the Tuticorin coast.

e­ xtensive intertidal flats. Near Puri the Black Pagoda, said to have been built on the shore, is now 16 km inland. Similar strandlines stand 5 m above sea level on the Baitarami delta. These features indicate emergence accompanied by the progradation of these deltas. Evidence of historical change is seen at Balasore, once a seaport but now 15 km inland. The escarpment behind Balasore looks like a sea cliff and isolated hills in this neighbourhood may be former islands: salt markets on the Brahmani River had to be shifted downstream. On the Ganges delta large gneiss pebbles (5–7 cm in diameter), some 160 m deep, probably indicate a low sea level during the Miocene. Subsidence in the Ganges delta during the Pleistocene and Holocene may be due to compaction and dewatering of sediment as well as isostatic sinking. The Bangladesh border crosses the Ganges delta.

4.  Andaman and Nicobar Islands The Andaman and Nicobar islands in the Bay of Bengal consist generally of Eocene to Pliocene conglomerates, sandstones and limestones injected by peridotite intrusions, and represent the emergent remnants of Tertiary fold mountains that once linked the Arakan Yomas of Burma to Sumatra. The indented coastline of these islands results from subsidence. They are located between 6° and 14° N, and it is possible that corals continued to exist here even during the Pleistocene glaciation, contributing to the formation of the uppermost coralline limestones. On Havelock and other islands of the Andaman and Nicobar

group, a dead forest standing in salt water is also an indication of subsidence. Raised coralline beaches on Great Nicobar Island, up to 15 m above mean sea level, indicate uplift. The 2004 tsunami was more devastating on the Andaman and Nicobar islands than on the mainland coast of India.

References Ahmad E (1972) Coastal geomorphology of India. Orient Longman, New Delhi Baba M, Thomas KV (1999) Geomorphology of the Indian coast. Ministry of Environment and Forests, New Delhi Chada RK, Latha G, Yeh H, Peterson C, Katada T (2005) The tsunami of the great Sumatra earthquake on 26 December 2004 – impact on the east coast of India. Curr Sci 88:1297–1301 Earattupuzha JJ, George V (1980) Shoreline changes on Kerala coast. In Geology and geomorphology of Kerala. Geological Society of India Special Publication 5:83–85 Hanamgond PT (2007) Morphodynamics of the beaches between Redi and Vengurla, Maharashtra, West Coast of India. Technical Report, Department of Science and Technology, Government of India. (ESS/23/VES/138/2001), 292 Karlekar SN (1993) Coastal geomorphology of Konkan. Aparna Publications, Pune Kunte PD, Wagle BG (2005) The beach ridges of India: A review. J Coast Res SI 42:174–183 Nayak GN (2005) Indian Ocean coasts – coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 554–557 Sambasiva Rao M, Vaidyanadhan R (1979) Morphology and evolution of the Godavari Delta, India. Zeitschrift für Geomorphologie 23:243–255 SAC (Space Application Centre) (1992) Coastal environment, Scientific note. Ahmedabad

17.3 Sri Lanka

Introduction The island of Sri Lanka has a coastline about 1,585 km long, increasing to 1,700 km when the Jaffna Lagoon is included. The climate is tropical, subject to the northeast (December to early March) and the southwest (late May to early October) monsoons: there is a wet zone in the southwest, where Colombo has a mean annual rainfall of 2,034 mm of which only 370 mm occurs between December and March, and a dry zone in the rest of the island. Winds rarely attain gale force except in gusts during the monsoons, but occasional cyclones (occurring about once in a decade) affect the northern coast. Waves are generated by the monsoons and local winds, and there is a southerly swell produced by storms in high southern latitudes and arriving across the Indian Ocean. This diminishes northward along the west coast, and is reduced by refraction along the east coast. Mean spring tide range is less than 1  m around Sri Lanka: Colombo and Trincomalee both have 0.6 m. Tsunamis are rare, but a major one occurred on Boxing Day 2004 as a result of an earthquake south of Sumatra. After an initial withdrawal, a series of waves up to 5 m high broke on the south and east coasts of Sri Lanka, causing extensive sea flooding and structural damage. On several sectors sand was washed onshore by these waves, and deposited on and behind pre-existing beaches. Sri Lanka is dominated by Pre-Cambrian schists and gneisses, except in the north and northwest, where there is Miocene limestone veneered with Quaternary gravels. The island is thought to have been tectonically stable in Holocene times, but more research is needed on emerged beaches and river terraces to determine whether ­warping and faulting occurred during the Pleistocene. Lowlands surround the Central Highlands, from which numerous rivers radiate, and the coast is low-lying, with few indentations. Much of the coastline is exposed to strong wave action, and long sandy beaches occupy arcuate bays between headlands of Pre-Cambrian bedrock (Swan 1982). On the

rest of the coastline, notably in the Miocene limestone zone in the north and northwest, wave energy is perennially low because of shallow seas and barriers, and the coastline is less regular in plan as the result of differential deposition. Beach sediment are predominantly terrigenous, quart­­ zose sands, with small and variable proportions of carbonate, notably close to coral reefs. The terrigenous sands are partly of fluvial origin, particularly on the east and west coasts, while the carbonate sands come from the sea floor or from disintegrating coral reefs. Seasonal alternations are typical: on the west coast the beaches are eroded during the southwest monsoon and largely restored, with the formation of a berm in the dry season. There has been a prevalence of erosion in recent decades on the beaches of Sri Lanka. Coastal dunes are present, mainly on sectors where rainfall is relatively low, with a pronounced dry season during which onshore winds blow sand in from the beach. Dunes occur on coasts where wave energy is vigorous, and also where it is low. Beach rock commonly outcrops on sandy shores where erosion has occurred, some exposures resulting from seasonal retrogression. Reefs of beach rock occur offshore, sometimes in a parallel, stepped series, probably in response to stillstands during the later stages of the Holocene marine transgression. Traces of emerged coastlines indicate higher Pleistocene sea levels up to 35 m above the present, and there are emerged Holocene shell beds along the south coast (Katupotha 1995). Sand derived mainly from the sea floor and swept in by ocean swell has been deposited in the form of spits and barriers that enclose coastal lagoons in former inlets and embayments. Sri Lanka has about 20 substantial coastal lagoons, with a water area totalling about 142,000 hectares. They include the much-modified Negombo Lagoon, north of Colombo, and a trio of lagoons on the south coast: Rekawa, which is tidal and mangrove-fringed; Kalametiya, where salinity has been reduced by the diversion of fresh water from a ricefield irrigation scheme; and Lunama, which is shallow and brackish (Bird 1987). Mangroves occur in sheltered lagoons and estuaries behind barrier beaches and spits where wave energy is

Edited version of a chapter by B. Swan in The World’s Coastline (1985: 749–759). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_17.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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considerable, and also in tidal flats open to the sea in the northwest where energy is low. Corals are widespread, and form fringing and some small barrier reefs: nearshore reefs occupy about 32 km of the coastline, mostly in the northwest (Rajasuriya and White 1995). Coral cover averages 50%, with Acropora, Montipora and Pocillopora spp. dominant. There has been much modification by human impacts locally, and some reefs are dead rock reefs with large boulders. Some fringing reefs decline into backing lagoons by way of a reef slope on which coral growth is sparse amid coralline rock outcrops with a patchy veneer of sand. Emerged coral reefs form low capes and headlands at many points around the island. The continental shelf extends across to India between the Gulf of Mannar and Palk Strait, but is narrow elsewhere, and generally divisible into an inner and an outer zone separated by the 55 m isobath. Beyond this the sea floor declines gradually until the continental slope begins at a depth of around 90 m. There is a net movement of sand northward, especially along the west coast of Sri Lanka, and there is probably an input into Palk Bay and the northeast from southern India. The coast at Colombo is beach-fringed, but a sandstone reef offshore excludes large waves and also reduces shoreward drifting of sand to the beaches, which have been eroding, exposing beach rock (Swan 1982). South from Colombo Pre-Cambrian bedrock outcrops intermittently along the coast. The outcrops are lateritised above the water table. Between the headlands are sandy beaches

backed by beach ridges, which in several sectors impound lagoons and swamps. Emerged beach ridges of Pleistocene age occur up to 8 m above sea level, and are up to 3 km wide at Colombo, but narrow southward to a few hundred metres. Between the headlands at Mt. Lavinia and Maggona the coastline is smooth and sandy, interrupted by the outflows of the Panadura and Kalu Gangas. South of Maggona, bays and headlands are backed by raised beaches, flood plains, swamps and lateritic ridges. Away from the protection of headlands, coast erosion is a hazard, with high wave energy along the southwest coast. Between Nam­ bimulla and Matara the problem is most severe because the sand supply is insufficient to compensate losses offshore and along the shore, and because there has been quarrying of coralline rock and removal of beach material. The Galle and Weligama bays are the only two large, relatively deep embayments on the southwestern coast of the island. At Galle a formerly eroding coastline has been stabilised by a sea wall (>Fig.  17.3.1), but wave reflection from this has dispersed the beach sand. Dondra Head, the southernmost point of the island, consists of low platforms of resistant granitic rock, capped by beach deposits. From Dondra to Tangalle numerous indentations make the coastline more scenic and varied than elsewhere in the island, with bays, promontories, headlands, cliffs and barrier beaches backed by swamps and lagoons. Rocky foreshores are exposed at low tide. Beach­ ­es  are discontinuous and coarse-textured, and have been

⊡⊡ Fig. 17.3.1 Sea wall at Galle, showing beach sand dispersal by reflected waves. (Courtesy Geostudies.)

Sri Lanka

derived partly from the erosion of weathered rock on headlands and foreshores, and partly from the sea floor. Little sand has been supplied by the few small rivers that reach this coast, most of which flow into lagoons. Waves are oblique to the coast in both monsoons, and the perennial southerly swell results in persistent wave attack upon the coast. East of Tangalle the coast declines to a broad lowland between headlands widely spaced, and long sandy barrier beaches fronting lagoons, estuarine deltas and swamps. Monsoon winds blow parallel to the coast, and the ocean swell, refracted to southeasterly across a broadening continental shelf, is largely constructive. Several rivers deliver sandy sediment to beaches on this part of the coast, but much of the sand comes in from the sea floor. Sand is washed into the mouths of inlets and embayments (>Fig. 17.3.2). From Ambalantota, west of Hambantota, to Sanga­ makande Point, the easternmost in the island, the coast consists of wide zeta-curved bays between bold headlands. Beaches are broad and dunes extensive. Although beach drifting alternates, the predominant longshore movement is eastward and northward. There is much sand in transit and marked accretion at the downdrift ends of beach compartments. There are signs that supply was even greater in the past. The hinterland consists of alluvial tracts and large lagoons, and inselbergs, tors, and other erosional remnants rising above plains cut in Pre-Cambrian rock thinly mantled with soil. Two linear submarine structures off the southeast coast are the Great and the Little Basses ridges: the name

⊡⊡ Fig. 17.3.2 Sand barrier deposited at the mouth of Lake Kalametiya. (Courtesy Geostudies.)

17.3

Basses is derived from the Portuguese Baixos, meaning a reef. Both are steep-sided, parallel to the coastline and rise from a bedrock and coralline sea floor that grades to a sandy bottom. The ridges are composed of sandstone cemented by calcareous material and probably mark submerged Pleistocene coastlines (Swan 1982). Alternatively, the ridges may be submerged coral reefs. On the east coast of Sri Lanka, between Sangamakande Point and Koddiyar Bay, coastal landforms of bedrock are rare. Instead, there are broad alluvial flood plains, with river distributaries that commence at elevations of 35 m up to 25 km inland. There are river terraces and several lagoons. Sand dunes diminish and disappear, because during the dry season when sand is mobilised, winds blow offshore. Barrier beaches are present, often with an inner and outer barrier. A series of large interconnecting lagoons collectively form the Batticaloa Lagoon, which has a northern outlet deflected by northward spit growth. Near Kalkudah and Periya Munai Point, the otherwise smooth depositional coast is interrupted by a series of headlands of emerged coral separated by bays backed by beach ridges, lagoons and low swampy terrain. Coral limestone has been intensively mined on this part of the coast. During the Late Quaternary marine transgression, the lower valley of the lower Mahaveli Ganga was submerged to form Koddiyar Bay. The Mahaveli has since built a Holocene delta, and sedimentation around the bay has produced a pattern of estuaries, deltas, lagoons and bayhead barrier beaches. The harbour of Trincomalee, with

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Sri Lanka

deep coves and inlets, receives no river sediment, and retains much of its earlier highly indented character. The bays at Koddiyar and Trincomalee open towards a major submarine canyon. Headlands and bays continue along the northeast coast of Sri Lanka. Between Trincomalee and Nayuru the coast is backed by raised beaches, lagoons, swampy alluvial flats and low residual hills with occasional narrow ridges trending northeast. Some ridges reach the sea as bold headlands. Other headlands are small, or are low capes of emerged coral, which succumb readily to northeast monsoon waves. Emerged beach deposits capped with relict dunes reappear, reddening the coastal sands, and between the mouths of the Yan Oya and Nayuru there are rich concentrations of ilmenite and rutile. From the mouth of the Nayuru lagoon northward the coast has extensively prograded, particularly at Mullaitivu, where there are large offshore banks of sand and the ocean swell is further weakened by refraction across a widening continental shelf so that it arrives from the east and northeast. To the NNW a wide barrier complex of beaches, scrubby dunes and small lagoons runs parallel to the coastline, and becomes the large tombolo that connects the limestone Jaffna peninsula to the mainland. The coastline swings westward as the limestone outcrop begins at Point Pedro. The northern coast of the Jaffna Peninsula consists of low (Fig. 17.4.1). The total length of coast, going around all the islands and up the estuaries, is estimated to be nearly 1,320 km. However, the width of the tidal channels, particularly in the Sundarbans, is impossible to measure with any precision, since the Bengal 1:250,000 scale maps are not very accurate. There is also the problem of how far inland the coastline should be considered to extend; tidal estuaries in places reach as far as 130 km inland, and storm surges may drive the Bay of Bengal waters up to 160 km inland. The climate is tropical, with heavy rainfall during the monsoon (June–October). Cyclones occur, particularly before and toward the end of the monsoon. Chittagong has a mean monthly temperature of 19°C in January rising  to 27.2°C in July, with an average annual rainfall of 2,831 mm. The Bangladesh coast can be divided into four parts, each with its particular characteristics. In the west, partly in India but mostly in Bangladesh, is the old delta called the Sundarbans. Next, to the east, lies a section where the Sundarbans have been cleared and there is considerable farming. The third division is the delta of the Meghna, the easternmost branch of the Ganges, and the fourth is the Chittagong coast from the Feni River to the Naf River at the Burma border, a distance of about 274 km. Except for the Chittagong section, the coastline is low, swampy, and rapidly changing. It comprises a large alluvial basin floored with Quaternary sediment deposited by two vast river systems, the Ganges and the Brahmaputra, with their numerous distributaries (Morgan and McIntire 1956). Discharge and sedimentary load figures are generally unavailable, but it can be said that the maximum flow of the Ganges River is 1.5 million cu. feet/s and the Brahmaputra some 2 million CCF/s. The geomorphological history of this huge deltaic plain is one of rivers that have continually shifted their courses back and forth as they fill in a large geosynclinal trough. In more recent times, the Ganges has

shifted to the east, from a main outlet along the western margins (the Bhagirathi–Hooghly in India) to the present main course, the Padma–Meghna; streams such as the Kobadak, Pursur, Madhumati, and Tetulia represent intermediate, but not necessarily successive, positions of the main channel. There are several possible reasons for this shift, but the tectonic subsidence theory advanced by Morgan and McIntire (1956) appears to be the most plausible. They suggested deep faulting or structural subsidence along a trough running north-northwest from the active delta to the junction of the Ganges and the Brahmaputra, southeast of Padma. Their ideas are based on the apparent structural trends of the rivers and on the faulting that occurs along the western edge of the Madhupur Jungle. Subsidence is taking place in the Ganges–Brahmaputra delta (Alam 1996). Whether any parts of the Bangladesh coast are prograding is a matter of debate. There is certainly erosion locally. There seems to be little difference between the coastline shown on Rennell’s map of 1770 and that of modern surveys, particularly in the western Sundarban region. However, the presence of old beach ridges in the western swamps points to earlier Holocene accretion. In the east, where large quantities of silt are brought down by the rivers in flood, stages in the process of land-building can be clearly seen from the air (Umitsu 1997). But no large delta appears to be extending into the Bay of Bengal, probably because of  subsidence and a fairly narrow continental shelf. The rate of subsidence has evidently more than kept pace with the rate of sedimentary deposition (Morgan and McIntire 1956). Tides along the Bangladesh coast are semidiurnal. Tidal heights vary with the season, being lower in February and March and higher during the monsoon, July through November. The mean spring tide range at the Pusur River mouth in the Sundarbans is 2.4 m, but tides here attain 7.0 m during monsoon storms (Sailing Directions 1978). At Chittagong mean spring tide range is 3.5 m and at Cox’s Bazar 3.0 m. Surface currents across the head of the Bay of Bengal are developed by the monsoon winds, and can be quite dangerous to small craft venturing into the bay.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_17.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

17.4

Bangladesh

⊡⊡Fig. 17.4.1 Predominant coastal landforms, Bangladesh.

KHULNA

BARASAL DISTRICT

R.

Estuarine Islands Mudflats and Shoals

R.

L

DELT

A

22°N.

NS

ARBA

SUND

CHITTAGONG

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Sub

BAY OF BENGAL

MAISKHAL IS. COX’S BAZAR

d s re

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m 305 etres Swatc metres h of n o gro un

et m 76

The Indian border crosses the Sundarbans, a heavily wooded mangrove and nipa palm swamp that extends from the Hooghly River in India eastward to the Tetulia River, a distance of about 280 km, of which nearly 193 km is in Bangladesh. The area covers about 6,000 sq km and represents the older deltaic plain of the Ganges River, with a seaward slope of about 0.076 m per mile. The largely uninhabited forest formerly extended northward to Khulna and eastward across the Barisal and Patuakha1i districts of Bangladesh, but with the encroachment of agricultural land and extensive cutting of the forests, the true mangrove and nipa jungles (>Fig.  17.4.2) are now largely confined to the southern part of Khulna district. The seaward fringe consists of swamp forests fronted by mudflats exposed at low tide (>Fig. 17.4.3). The region is interlaced with a checkerboard pattern of tidal estuaries and tidal rivers. Many of these interconnecting channels are navigable. Interstream areas are

BURMA

T

2.  The Bangladesh Coastline

AS

80 km

CLIFFS AND BLUFFS FRONTED BY NARROW BEACHES

CO

PREDOMINANT COASTAL LANDFORMS

N

Alluvial Plains

KUTUBDIA IS.

BANGLADESH

21°N.

.

li R

phu

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Ka

CHITTAGONG

TIDA

SANDWIP IS.

NA GH A ME ELT D

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TU DI AK ST HA RI L I CT

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Non-Saline Tidal Delta

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CL COEAR AS ED T

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INDIA R.

eg

SALINE DISTRICT

CALCUTTA

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INDIA

Pa dm a

M

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23°N

um uh ad M

MORIBUND DELTA

TROPIC OF CANCER

92°E.

91°E.

DACCA or

Feni R.

90°E. ng es

89°E.

Ga

88°E.

B

1078

SAND BEACHES AND SHOALS SALT MARSHES AND SWAMPS SUNDARBAN FORESTS CHITTAGONG HILL TRACKS

Naf R.

shallow, saucerlike swamp depressions. Off the coast are a number of low, flat marshy islands. Navigation in the channels and estuaries is hazardous, especially during the “northeasters” in March, April, and May and the Bay of Bengal cyclonic storms from June to October. Tidal mangrove and sundri (Heritiera littoralis) forests predominate. The mangrove forest, mainly species of Rhizophora, with some Avicennia, is found along the waterways, the banks of which are composed of soft mud and clay that is under water at every high tide (>Fig. 17.4.3). Many of the creeks and channels are fringed by nipa palms (Phoenix paludosa). Higher up is found the sundri forest, where the land is flooded with moderately brackish water at every high tide. The sundri tree reaches a height of about 16 m, and its red wood is of excellent quality. Close to the coast the high forest changes into lower species, separated in places by low sand dunes. The Sundarbans are the habitat of cobras, pythons, crocodiles, leopards, small deer, and the famous Bengal tiger, whose numbers are now being threatened by man.

Bangladesh

17.4

⊡⊡Fig. 17.4.2 Nipa palms and mangroves in the Sundarbans.

⊡⊡Fig. 17.4.3 Mudflats and Rhizophora on the Tetulio River delta.

East of the Tetulia River, in the Barisal and Patuakhali districts of Bangladesh, over a distance of some 68 km, the Sundarban forest has been cleared (> Fig. 17.4.4), and there is extensive farming (mainly paddy rice) on the drier and less saline interfluves of this part of the deltaic plain. Here population densities reach more than 2,900 per sq km. Along the coast lie scattered, partially cut mangrove forests. The Ganges, Brahmaputra, and Meghna rivers all join into one main trunk stream in central Bangladesh before

entering the Bay of Bengal through a single main channel called the Meghna River. Where it reaches the Bay of Bengal, it divides and spreads out around a series of extensive shoals called the Meghna Flats. These largely barren mud and sand shoals, at least ten of them of considerable size, extend from the Tetulia River on the west to the Feni River on the east, a distance of about 100 km. Wide, shallow channels separate the shoals, shifting so often that navigation is dangerous for large boats, especially during

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Bangladesh

⊡⊡ Fig. 17.4.4 Cleared mangrove forest land in agricultural use, Patuakhali district, showing coastline erosion.

monsoon floods and severe cyclones. Only vessels with a draft of 3 m or less can proceed up the channels. The mud and sand shoals, with depths up to 5.5 m, extend southward for some 70 km to the Karnaphuli River. This is a drowned coastal region with no true delta building out into the Bay of Bengal (Singh 2002). During storms, waves cover these shoals and, in contrast to the Sundarbans, little vegetation survives on them. The coast of the Chittagong district extends 274 km from the Feni River in the north to the Naf River on the Burmese boundary. The small Feni River empties into Sandwip Channel about 13 km above the northern end of Sandwip Island. Two other large islands, Kutubodia and Maiskhal, are found along this stretch of coast, as are several smaller islands and shoals. The coastal plain varies greatly in width here. Near the Karnaphuli River, it is from 48 to 64 km wide, but where the low Tertiary hills and cliffs rise immediately inland, it is virtually nonexistent. The Karnaphuli River, with its headwaters in Assam, is short but carries a large volume of water during the monsoon flood season. The land bordering the mouth of this river is low and flat, and the entrance would be difficult to  find by ship but for the Norman’s Point Lighthouse.

S­ and-covered mudflats, dry at low tide, extend about 1 km seaward of the mouth. Several small beaches and broad sand flats are found along the coast between the headlands, but in the south the coastal plain dwindles away before the uplifted and deeply eroded Chittagong Hills, which inland reach an elevation of 609 m. These hills are of stratified Tertiary marine and freshwater limestone and shale. At Cox’s Bazar, there is a beach more than 152 m wide, but south of the town steep cliffs 9–15 m high front directly onto a narrow shore.

References Alam M (1996) Subsidence of the Ganges-Brahmaputra delta of Bangladesh and associated drainage. Sedimentation and salinity problems, vol 2. Kluwer Academic Publications, the Netherlands, pp 169–192 Morgan JP, McIntire WG (1956) Quaternary geology of the Bengal Basin. Coastal studies Institute. Technical Report 9. Louisiana State University, Baton Rouge, LA Singh OP (2002) Predictability of sea level in the Meghna estuary of Bangladesh. Water Energy Int (58):13–19 Umitsu M (1997) Landforms and floods in the Ganges Delta and coastal lowlands of Bangladesh. Mar Geod 20:77–87

17.5 Burma (Myanmar)

Kyaw Saw Lynn

1.  Introduction The coastline of Burma (Myanmar) is about 2,300  km long, and consists of three distinct parts (>Fig. 17.5.1): the western Arakan (Rakhine) coast, the southern Irrawaddy (Ayeyarwady) deltaic coast, and the north-south Tenasserim (Tanintharyi) strip. The Arakan coast, from the Bangladesh border south to the Cape Negrais peninsula, runs parallel to mountain ranges of strongly folded Mesozoic and Tertiary rocks, dominated by ridges and valleys that follow the geological strike. East of Cape Negrais the deltaic plains of Irrawaddy border the Gulf of  Martaban (Mottama), and south from Moulmein the coast, which has a predominant north-south strike, is again generally steep, bordering the Tenasserim Ranges of mainly Upper Palaeozoic rocks. In a humid tropical setting, with annual rainfall exceeding 5,000 mm in many parts of the coast and only a brief dry winter season, rock formations are deeply weathered, and coastal slopes carry evergreen rainforest or bamboo thickets. Oceans waves arriving in coastal waters are generally weak. However, they are reinforced by the southwesterly monsoon in late May to early October to produce moderate wave action on sandy beaches and rocky shores with minor cliffing. Tropical cyclones are rare in this part of the Bay of Bengal. Tidal ranges are typically about 2.0 m as at the mouth of the Bassein River, increasing to more than 5.5 m in the estuarine distributaries of Irrawaddy, and to more than 6 m at the head of the Gulf of Martaban, where the Sittang River, opening into an estuarine gulf, is invaded by a tidal bore at spring tide. The  range at Moulmein is 3.6 m and at Mergui, to the south, is 5.3 m.

2.  The Burmese (Myanmar) Coastline South of the Bangladesh boundary, which runs down the Naff Valley and along a marine inlet, the northwest Arakan coast is of Dalmatian type, with elongated steep-sided peninsulas and islands separating marine straits. The coast mainly exists as reefs of origin; primarily, coral and ­tectonic

movements have played an important role in the development of coastal landforms found along the Arakan coast. The extensive level lowlands in the Arakan coast are mostly made up of alluvium, fed by rivers that drain the northwestern stretch of the country. The higher grounds, on the other hand, are composed of sedimentary rocks formed during the Tertiary period, and generally rise to the east to more than 1,500 m above sea level. The characteristics of the Dalmatian coast are particularly evident along the stretch from Naff to Sandoway (Thandwe) River. Here, offshore islands which are aligned in a northwestly direction along with the Arakan mountain range (Rakhine Yoma), parallel the coastline. These offshore islands are connected to one another and with the sea by means of an extensive network of vast estuaries and creeks. Among the offshore islands, which are tectonically associated with the Arakan mountain range, Ramree Island is the largest, while the Cheduba Island (Man-aung Island) is the second largest. There are intersecting estuarine channels behind Ramree Island. Mangroves are extensive, particularly around river mouths, and coral reefs are patchy, mostly on the outer coastline. Emerged coastal features at various levels indicate a history of tectonic movements, and the 1762 Chittagong earthquake is said to have caused actual emergence around high Cheduba Island, which has at least three well-developed raised coastlines. Kyakpyu is a town at the northern end of Ramree Island, which consists of former islands, now hills, rising above emerged alluvial plains. Raised beaches and old sea cliffs are preserved on the west coast, particularly at Minbyin (Pascoe 1962). From Sandoway (Thandwe) to Cape Negrais, the features of the coastline change. The coast mainly takes the form of a rugged and rocky stretch, with a few available places of refuge for vessels in distress. Towards the south the coast is dominated by a series of precipitous headlands, which are offshoots from the Arakan Yoma and separated by sandy beaches. The magnificence of this coastline is encapsulated in the form of Ngapali beach, close to Sandoway, the most popular pioneer resort destination in the country. The Arakan coast, despite its rocky topography, has another leading seaside resort, known as Kanthayar Beach ­(literally

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_17.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

17.5

Burma (Myanmar)

⊡⊡ Fig. 17.5.1 The Burmese coastline. (Courtesy Geostudies.)

R

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Folded Sedimentary Rocks Upper Palaeozoic Igneous and Sedimentary Rocks Alluvium 0

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Burma (Myanmar)

meaning tranquil coastline), near Gwa. Overall, the Arakan coast is indented by spacious inlets and lagoons with scattered sandy bays. Generally, soil erosion in Arakan coastal catchments has augmented river loads; as a result, there are extensive tidal mudflats in sheltered areas, as in Gwa Bay, Danson Bay, and Ngayak Bay along the southern stretch of Arakan coast. Sand spits, as at Cypress Point and the mouth of Sandoway River, have grown to deflect river mouths and protect mangrove areas. Mud volcanoes are a feature of these coastal waters, forming temporary islands up to 4 m high, which are planed off by waves to form shoals. There are good examples in Cheduba Strait, where they are a hazard to navigation. They are evidently associated with submarine volcanic activity where the Indian continental plate is under-thrusting the Asian plate. Foul Island, to the south, is a thickly wooded volcanic island that has erupted in recent decades. The steep, embayed coast continues southward to Cape Negrais, at the end of a ridge on the western side of the Irrawaddy delta. Beyond Pagoda Point (Mawtinsoon) the coastline turns east, and its character and appearance change, the bluffs and rocks giving place to flat and sandy beaches with narrow grass-covered plains running along its margin. Numerous mud and sand banks, many dry at low tide, stretch out seaward, rendering the whole coast unapproachable by vessels of large size, within 9 or 10 miles, except in a few places where channels are kept open by the streams or rivers. The Arakan coastline has been shaped by Late Qua­ ternary marine submergence. There are branching estuaries, drowned valleys and, in the east of the Ramree Island, an extensive lagoon with numerous islets and mangrove swamps. There is also evidence of coastal emergence. On the Baronga and Ramree islands in Akyab (Sittwe) area, a number of well-defined emerged bars stand at heights of  many feet above sea level, and run in successive rows  inland. On Round Island there is a table-like cliff rising to an emerged marine terrace at a height of about 12 m. Terraces found on the Terrible, Foul, Flat, and Cheduba islands on the Arakan coast also indicate recent upheaval. Off the mouths of the Kaladan River and Laymyo River on the Arakan coast there are three narrow, elongated, parallel islands, called the Baronga Islands, which are really three detached ranges of low hills running southward into the sea. Their direction is north-northwest to south-southeast, and the rocks are evidently a continuation of those in Ramree and the Cheduba further south. A  small amount of fair quality petroleum has been ob­­ tained from these islands. Further south, there are two

17.5

large islands, Ramree and Cheduba, and several small islets, which may be conveniently grouped together as the Ramree Group. The Ramree Group is evidently an emerged archipelago, consisting of an extensive silty plain. The islands are generally swampy, and afford excellent facilities for growing rice. Most of the creeks on the islands are tidal and contain brackish water. Hills rise up to 150 m and are covered in thick jungle. Emerged beaches of shingle and shells, and old sea cliffs, are to be seen along the west coast. The Ramree Group has produced oil on a small scale. Most of the oil in the Minbyin field is taken from depths between 75 and 90 m, although some of the wells are nearly 150 m deep. The Cheduba, to the south of the Ramree Island, is a fertile, vegetated island of moderate height and irregular outline. It contains many irregular, low, undulating hills, varying in height from 15 to 150 m, enclosing several higher detached mounds with steep, well-wooded slopes, the highest of which, near the southern part of the island, rises nearly 420 m. The Cheduba has a large and locally celebrated mud volcano in the northwestern corner. It has a crater surrounded for about 8 ha with hardened mud, which is cracked in many places, with bubbling mud in the crevices. Petroleum in small quantities has been found at ten localities on the island. In the Ramree Group there are also several smaller islands, some of which are quite recent in origin and formed only in the early years of the present century. They  include Round Island (Toquequong), Flat Island (Rekeong), Amherst Island (Jergo), Volcanic Island, Foul Island (Nanthakyun), False Island, and Beacon Island. Volcanic Island (1° 90' 16" N, 93° 24' 20" E) is a mud volcano formed on 15 December 1906, 12 miles distant and 36° West of North from Beacon Island, off Cheduba. The maximum length of the island is 280 m and the maximum breadth 200 m, its summit being 5.8 m above high water. The island is composed wholly of greyish-brown mud, with a few angular fragments of rocks of various kinds thrown up with the mud. The Irrawaddy River, the longest river in Burma, flows into the Andaman Sea through nine major distributaries, forming the Irrawaddy Delta, which covers an area of 35,000 sq km, one-sixth of which is intertidal, submerging at high spring tides. The Irrawaddy Delta extends from theRangoon (Yangon) River on the east to the Bassein (Pathein) River on the west. The delta is mainly composed of silt and sand deposited by the Irrawaddy and its tributaries, which drain into a major tectonic trough that runs north to south. The seaward margin of the delta is vegetated with mangroves, forming forests up to 30 m high,

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17.5

Burma (Myanmar)

and associated dani palms, spread out on to accreting tidal flats, especially on the seaward margin in the lee of wavebuilt sand bars. Large mats of waterweed are carried downstream by the Irrawaddy during the wet season, and these eventually become incorporated in accreting sediment. Locally, the river channels have meandered to undercut and cliff the margins of residual hills of older sandy alluvium. Rangoon, sited where a spur of the Pegu Yoma runs out into the deltaic plains, is a major port on a navigable estuary. In addition, extensive flooding occurs in the wet season, and high natural levees have been built by silt deposition alongside the river channels. The Irrawaddy Delta and the Gulf of Martaban (Mottama) are believed to originally have been formed as a part of the subsiding geosynclinal tract, which received the Pegu and Irrawaddian sediment. Subsidence was interrupted by the broad corrugation of Pegu Yoma. Rivers ancestral to the Irrawaddy River supplied sediment that filled the Tertiary Gulf. Infilling of a subsiding Burmese Tertiary Gulf by the southward encroachment of deltaic sediment supplied by the tributaries of Irrawaddy must have taken place, much as the present Irrawaddy Delta is encroaching upon the modern Andaman Sea. It is thought that the Andaman Sea started forming by rift movements initiated during the late Miocene. The geology of the Irrawaddy Delta can be examined in northern and southern sections. Two types of Irrawad­ dian sediment are found: continental and marine. The northern area is occupied by sandstones that are primarily light-coloured, medium- to coarse-grained, angular to sub-angular, poorly to moderately well cemented, generally non-calcareous, and rarely ripple-marked. Reworked shales are also frequently found, while conglomerates and  quartz pebbles are either disseminated or in bands. Ferruginous and calcareous concretions of secondary origin are common, usually resembling large cannon balls, sometimes disc-shaped. In the southern area the sedimentary formations are similar, except for the presence of more argillaceous beds near the base of the formation. Fossil wood is rarely found, probably due to the increased marine influence in the south. The Irrawaddy Delta is generally advancing. Observa­ tions carried out by the Myanmar Naval Hydrographic Department in 1950 and 1994 show that during this period new islands were formed, particularly between the mouths of Eyar River and Toe River. These included Kaing Thaung Island, at the mouth of the Eyar; Kyun Thayar, Khattar, Gayet Kyee, and Ngaman Thaung islands at the mouth of the Bogalay River. Myet Sein Island is located at the mouth of the Toe River.

Sedimentation and delta building in the Irrawaddy Delta has been continuing for thousands of years. Enor­ mous quantities of sediment have been brought down to the Katpali Sea and Gulf of Martaban by the Irrawaddy, and long-term deposition by the Irrawaddy and Salween rivers has contributed to the seaward expansion of the delta and deltaic shelf. The delta grows along the coast by the coarser sediment discharge, silting the shallow sea. Progradation occurs at up to 60 m per year on this deltaic coastline, even though much silty sediment delivered by the rivers is swept eastward into the Gulf of Martaban (Volker 1966). The Irrawaddy Delta advanced seaward at an average of 5.2 km between 1929 and 1953, and the mangrove swamp areas advanced up to 8 km during this period. The Tenasserim coastline stretches for about 1,360 km, from Bilin to southernmost Burma at Victoria Point. This southern coastline resembles the western (Arakan) coastline, with its numerous offshore islands and irregular outline. The Tenasserim coastal area is bordered by the Kingdom of Thailand to the east and south, and by the Andaman Sea to the west. The north-south orientation of the coastline is greatly influenced by tectonic alignments in southern Burma. The Tenasserim mountain  range (Tanintharyi Yoma) is regarded as the southern extension of the northeastern Shan Highlands of the country. Like the Shan upland, the Yoma is composed of Archaean Pa­­ laeozoic and Mesozoic rocks, and can be considered part of Indo-Malaysian mountain system. In the northernmost part of Tenasserim the Sittang River flows into the head of the macrotidal Gulf of Martaban, the east coast of which is an embayed lowland extending south to Moulmein and the Tenasserim Coast. South of Moulmein the coast borders steeply rising country, but there are a number of sectors of alluvial plain. The Salween River opens into a funnel-shaped estuary bordered by plains from which steep karstic hills of limestone rise. Moulmein stands on its eastern shore, protected from the Andaman Sea by hilly Bilugyun Island. Spurs of the hill country extend to the coast, ending in rocky promontories; the intervening valleys open to mangrove-fringed creeks and estuaries, often with bordering sand spits and outlets, usually at the southern end. Mining has been extensive in the Paleozoic rocks of the steep hinterland, and extraction of minerals and precious stones from valley floors and terrace alluvium has modified runoff and sediment yield from many rivers, but the effects of these activities on the coastal landforms have not been documented. Many estuarine lagoons in bays along the coast have been rapidly silted and converted into mangrove

Burma (Myanmar)

swamps and saline marshes. Coastal sand deposits include heavy minerals: notably ilmenite and monazite. Some beaches are backed by parabolic dunes blown inland by the southwesterly monsoon winds. Amhert (Kyaikkami) stands on an alluvial lobe south of Bilugyn Island. To the south the coastal plain is interrupted by ridges, breached by Heinz Bay, and continuing south to Tavoy Point, backed by the long Tavoy estuary. The ridge continues in Mali Kyun (Tavoy Island) and passes through Kadan Kyun, off Mergui. In the middle section of the Tenasserim coastline, near the Javoy (Dawei) River, are numerous sand spits and beaches, Maungmagan beach being one of the best known in Burma. The southern coastline has many bays and headlands, and is fringed by numerous islands. The Mergui (Myeik) Archipelago is a group of more than 800 islands along the southern coastline of Burma. The inner islands are vegetated with mangroves, but the outer ones can rise steeply from the sea to a height of a few hundred metres. The Tenasserim coastline, like the Arakan, displays evidence of emergence and submergence. Emergence has occurred along the coast from Bilin to Javoy, where a plain separating the Andaman Sea from the chain of ridges in the Bilin-Thaton-Martaban area and areas further south undoubtedly represents a continental shelf that has re­­ cently emerged from the sea. The hills sloping in the direction of these plains show evidence of recent marine denudation, and there are wave-worn rocks north of Martaban and wave-cut notches north of Mudon and at Kyaiktalone. Where the hills rise behind the coastal plain there are emerged islands, peninsulas, and bays. A wide, straight, emerged sandy barrier, devoid of vegetation, extends north and south of Heinze Chaung. Between this barrier and the sea there is a series of parallel vegetated beach ridges. Former cliffs are also found behind the flat surface of the emerged shelf south of Maungmagan. Along the stretch of the Tenasserim coast south of Javoy there is evidence of submergence, as in the lower part of the estuary of Javoy River, which has not yet formed a delta or filled its estuary in spite of an abundant sediment load. Similarly, the mouths of the Great Tenasserim River, Lenya River, and Parchan River remain unfilled, despite the enormous discharge brought down by these rivers. The intricate mainland coast has mangrove swamps between steep-sided headlands, and a mountainous ridge

17.5

declining to Kawthaung, backed by the Lenya estuary along the Thai border. The Mergui Archipelago, stretching north to south along the Tenasserim coastline, consists of some 804 hilly and forested islands. The islands vary in size, from a mere outcrop of rock to a series of pinnacles to the largest King Island, which has an area of 440 sq km. Most of the islands are dominated by luxuriant vegetation and primary forest, housing some of the last jungle cats and other large mammals to be found in Southeast Asia. Most of the islands are fringed by mangrove swamps, but occasionally have yellow sand and gravel beaches. Historically, the archipelago was an important trading centre between the East and the West, particularly in the eighteenth century. There are shallow, inshore fringing reefs, offshore fringing reefs, pinnacles and small rocky islands, and banks that rise up from depths of over 300 m. Many of the small islands in the Mergui Archipelago are composed of metamorphic rocks, generally referred to as part of the Mergui series. Some of the islands are undoubtedly the summits of a submerged hilly area of granitic rocks. The islands of Mergui Archipelago appear to follow four tectonic lines. The innermost line starts from islands a little south of Palauk River and run southward to the islands near Karthuri. The second line of islands breaks off from the mainland at Tavoy Point, and  includes Tavoy Island, Domel Island, Lampi Island, King Island, and St. Matthew Island. The highest point is 767 m on King Island. A subsidiary line of islands extends from King Island through the islands of Sellor, Forbes, and Kisseraing, joining the mainland south of Campbell Islands. The third line starts from a point near Maungmagan and passes through the islands of Ross, Maria, and Clara to end near the Sayer Islands in the extreme south, where the sea deepens to the west. The outermost line passes through the islands of Cabusa, Tanintharyi, and Hayes.

References Pascoe EH (1962) A manual of the geology of India and Burma, Vol 3. Geological Survey of India, Delhi Volker A (1966) The deltaic area of the Irrawaddy River in Burma. In: Scientific problems of the humid tropical zone deltas and their implications. UNESCO, Paris, pp 373–379

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18.0 Indian Ocean Islands – Editorial Introduction

Introduction The Indian Ocean includes some islands that are fragments of continental structures, such as Madagascar and the Seychelles, others that are of volcanic origin, such as Rodrigues, and many that are coralline, such as the emerged atolls of the Chagos Archipelago and the cays of the Maldives (Wafer et al. 2005). >  Madagascar is an island of continental metamorphic and sedimentary rocks, with a dominant north-south trend seen in a major escarpment that overlooks the east coast and rises to a range that forms a divide between longer west-flowing rivers and shorter rivers that descend to the east coast. Southeasterly ocean swell and waves generated by southeast trade winds break on long sandy barrier beaches on the east coast, backed by lagoons. Ridges run out to the irregular northwest coast where coral reefs are extensive, and there is evidence of tectonic uplift. The west coast borders a broad coastal plain of Quaternary sediment. Mean spring tide ranges are between 3 and 4 m on the west coast bordering the Mozambique Channel, but less than 1 m on the east coast.

 Other >Indian Ocean Islands include the Comores archipelago in the north of the Mozambique Channel, which consists of volcanic islands that have been cliffed, bordered by beaches of volcanic and coralline sand and gravel, with fringing coral reefs. Reunion, Mauritius, and Rodrigues are other islands of volcanic origin. In the Seychelles there are granitic islands, and in the eastern Indian Ocean the Andaman and Nicobar Islands are the partly submerged southward continuation of a Burmese mountain range. Some of the other Indian Ocean islands are uplifted coral reefs. They include Aldabra and Christmas Island. Low, sandy islands (cays) of coralline sand and gravel derived from reefs and deposited on a reef platform occur in the Maldives, the Laccadives, and the Amirantes (Seychelles).

Reference Wafar M, Wafar S, Yennavar P (2005) Indian Ocean Islands, coastal ecology and geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, pp 557–564

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_18.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

18.1 Madagascar

Jean-Michel Lebigre*

1.  Introduction Madagascar has a 6,000-km coastline (>Fig. 18.1.1), and is 1,600 km from north to south, with a maximum width of 600 km. It lies only 300 km off the east coast of Africa, from which it is separated by the Mozambique Channel, more than 4,000 m deep. According to geologists, the island was attached to Africa about 165 million years ago. As the continents drifted apart, the island of Madagascar broke away from Africa, and its fauna and flora have evolved along separate lines ever since. The island is markedly asymmetrical in both its topo­ graphy (with a gradual slope toward the Mozambique Channel and a sharper descent toward the Indian Ocean) and its geology. Madagascar includes large sedimentary basins to the west, a crystalline shelf in the central and eastern highlands, and only a narrow, discontinuous sedimentary fringe on the eastern side. As a result of the topographic asymmetry, the largest rivers drain to the west coast, the eastern side having numerous shorter streams. The continental shelf is wide off the west coast and in the far south, but narrow off the east and northeast coasts. There is also climatic contrast. The eastern slopes are exposed to the southeasterly trade winds and thus very wet (rainfall 2,000–4,000 mm annually), while the western slopes are drier, and the southwest semi-arid: Toliara (formerly Tuléar) has a mean annual rainfall of only 350 mm. Madagascar has many fine coral fringing reefs and barrier reefs, wide and deep lagoons, coral banks with cays, and a rare drowned barrier reef off the northwest coast. The best developed coral reefs are found in the southwest, some fringing, others of the barrier type. In the north fringing reefs are well developed in the Antsiranana region, at Nosy Be, on the Ampasindava peninsula, and locally within the Bay of Narinda. Between the delta of the Mangoky and Cape Saint-André there are numerous reef banks bearing cays. On the east coast there are also discontinuous fringing reefs from the Masoala Peninsula to the Tamatave sector. Farther south the reefs disappear along more than 1,000 km of coastline.

Mangroves and salt flats cover about 425,000 ha, with seven mangrove genera (Rhizophora, Bruguiera, Ceriops, Avicennia, Sonneratia, Xylocarpus, and Lum­nitzera). The most widespread mangroves occur on the northwestern and western coasts, while on the east coast south of Toamasina and on the south coast, they are rather rare and restricted. Marine deposits dating from the last interglacial have been found and dated on various parts of Madagascar. They generally stand above present sea level, but in the region of Nosy Be and around Bas Sambirano in the far northwest the interglacial coastline has been downwarped below present sea level by subsidence between faults trending north-northeast to south-southwest.

2.  Coast of Madagascar The coastal landforms of Madagascar are shown in > Fig.  18.1.2. The east coast, exposed to southeast trade winds, has a high rainfall (up to 4,000 mm/year) whereas the west coast is drier and the southwest semiarid: Toliara has an annual rainfall of 350 mm. From rugged basaltic Cap de Ambre (Cape Bobaomby), the northernmost point, the coast trends southwest to Cap St André (Tanjona Vilanandro). This is the boldest coast in Madagascar. There are several rias and embayments, formed where the Flandrian marine transgression invaded a landscape of basaltic rocks on the coast of the Montagne d’Ambre, clayey hills in the Ampasindava peninsula (up to 730 m high), and karstic topography on the limestones of the Narinda Peninsula. In the northwest fringing coral reefs are well developed in the Cape Bobaomby region, at Nosy Be, on the Ampasindava Peninsula, and locally within the Bay of Narinda. Off Mahajanga there is a drowned barrier reef with generally reduced coralline life, formed by sandy banks beneath about 20 m of sea water, located on the outer edge of the continental shelf and bordered by deep sea (about 2,000 m). There are emerged coral reefs up to 25 m

*Edited version of chapter 18.1 (Madagascar) in The World’s Coasts: Online (2003) by Jean-Michel Lebigre (University of New Caledonia, New Caledonia). All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_18.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Madagascar

⊡⊡Fig. 18.1.1 Madagascar.

18.1

Madagascar

⊡⊡Fig. 18.1.2 Coastal landforms of Madagascar. (Courtesy R. Battistini.)

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Madagascar

⊡⊡Fig. 18.1.3 Cliffs in the Pliocene sandy clays north of Mahajanga.

⊡⊡Fig. 18.1.4 Basaltic cliffs on the coast of Nosy Be.

above sea level on the northern peninsula, and at 3 m out on the Iles Glorieuses, northwest of Cap d’Ambre. The finest examples of ria morphology are the Loza Ria and the large bay of Narinda. Those at the mouths of the Mahajamba and Betsiboka (also named Bay of Bombetoka) result from the submergence of cuesta relief, the Flandrian transgression having led to penetration of the sea through consequent gaps cut by the rivers, ­widening in the ­subsequent depressions behind the cuestas to form broad bays with narrow entrances. Three rivers have formed extensive deltas: the Northern Mahavavy near the Ankarana karstic massif, the Sambi­ rano, and the Southern Mahavavy west of Mahajunga, the major town of the northwest coast. This area has the most

extensive mangroves in Madagascar, within the larger bays and along delta fronts. The Sofia River flows into the Bay of Mahajamba, which has extensive mangrove swamps, mudflats, and tidal inlets along its southern shore. The sea is stained red by sediment derived from lateritic soils in the river catchments. The braided channels of the Betsiboka River discharge an extensive plume of red sediment. Although much of the coast is low-lying the sea has cut high cliffs in Pliocene sandy clays near Mahajunga (>Fig. 18.1.3). There are several offshore islands, including Nosy Mitsio, Nosy Be, Nosy Radama, and Nosy Lava. Volcanic Nosy Be is the largest such island (293 sq km), with cliffs of volcanic rock (>Fig. 18.1.4).

Madagascar

18.1

⊡⊡Fig. 18.1.5 Aerial view of the southern part of the Mangoky delta: mangroves, salt flats, and sand ridges.

⊡⊡Fig. 18.1.6 Salt flat at Ankiamena on the delta of the Tsiribihina.

On the west coast of Madagascar, between Cap St André and the Mangoky Delta (>Fig. 18.1.5), the coastline is simpler, with no major indentations. It is low-lying, with  extensive alluvial plains, including several deltas. The coastal landscape is one of sandy barriers, behind which are brackish and saline lagoons, extensive mangrove swamps, and wide salt flats (>Fig.  18.1.6). Some progradation has occurred on delta shores, but parts of the coastline are retreating as the result of marine erosion, such as around Maintirano and on the Manambolo Delta.

Since 1923 coastal erosion has been severe at Moron­ dava, near the mouth of the Morondava River. From 1987 to 2000 the coastline retreated 30 m, much of this during Cyclone Cynthia in 1991. Morondava has been protected by the building of a series of groynes at right angles to the coastline. Offshore, beside the deep Mozambique Channel, the continental shelf is wide, and there are numerous reef patches bearing cays. These include many small sandy islets off Maintirano and south of Morondava, notably Nosy Barren, Banc Pracel, and Nosy Andriamitaroka.

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18.1

Madagascar

South of Morondava the low coast continues past the Mangoky delta to Morombe. South of Morombe, at Andavadoaka, there are emerged reefs edged seaward by notches and visors. Further south is the small Baie des Assassins, and a dune-fringed coast extends to Manombo. Between Manombo and the mouth of the Mandrare is a semiarid area, with extensive Quaternary coastal dunes. Along 650 km of coastline there are only a few river mouths: the only rivers are the Manombo, Fiherenana, Onilahy, Linta, Menarandra, Manambovo, and Mandrare, all intermittent with the exception of the Onilahy, which flows all year round. After seasonal rains inland, streams occasionally supply the coastal lagoons with fresh water. Locally, in the Bay of Saint-Augustin (at the mouth of the Onilahy River), the sea has cut the high cliffs of Barn Hill in Eocene marine limestone. Mangrove areas are discontinuous (Lebigre 1997). They have colonised the shore south of Morombe (the Tsingilofilo Marsh), the Helodrano Fanemotra (formerly Baie des Assassins), and Saint-Vincent Bay. They are also  well developed on the Manombo and Fiherenana (>Fig.  18.1.7) deltas, in  the Onilahy estuary, along the Mangoro indentation, on the Linta Delta close to Androka, and behind the longshore spit in the Bevoalava Lagoon (>Fig. 18.1.5). Seagrass beds are extensive. There has been extensive sand deposition. The large Sarodrano sand spit has grown south of the provincial

capital, Toliara (Battistini 1995). On the nearby delta of the Fiheranana, barchans have advanced over saline flats. Active Holocene dunes border the delta (>Fig.  18.1.8), and in these are found eggshells of Aepyornis, a giant extinct bird. Sandy beach ridges and cheniers occur on salt flats and in mangrove areas. Long sections of sandy coastline at Mikea and Maha­ faly, north of Itampolo, are backed by active dune fields, dominated by parabolic dunes oriented in relation to the prevailing southerly and southwesterly winds. The older dunes, including the weathered reddish dunes of Tatsimian (early Quaternary) age and the pink or rose-coloured dunes of Karimbolian age (150,000–80,000 years bp), have been lithified into sandstone, and in places form ridge segments 10–15 km long, parallel to the coastline. Marine erosion has attacked these calcareous aeolian sandstones (dune calcarenites), forming cliffs a few metres high, rising to 200 m at Cap Sainte Marie, which is the southernmost point of Madagascar. There are small cliffs on the edge of the Tsingilofilo salt flat. There are fringing coral reefs on the Mikea and Mahafaly coasts, and barrier reefs from the Bay of Assassins to Morombe and the Grand Récif off the coast of Toliara, notably between Manombo and Belalanda. The continental shelf narrows to the south, and is trenched by a deep submarine canyon off the Onilahy River. Cliffs cut in dune calcarenite extend from Cap St. Marie eastward to Faux Cap, and dunes fringe the

⊡⊡Fig. 18.1.7 Mouth of the intermittent Fiherenana River after enlargement in 1989. The former waterway is recognisable in the photo by the sand line.

Madagascar

18.1

⊡⊡Fig. 18.1.8 Drifting Holocene dune on the Fiherenana delta.

steep  coast from Faux Cap to the Mandare River. Lac Anory, to the east, is a lagoon sealed off from the sea by a  barrier bearing drifting dunes. Between here and Tolanaro (formerly Fort-Dauphin) there is a short sector of indented coastline, with rocky capes corresponding with the extremities of the granitic Anosyenne Chain, separated by infilled rias, inlets, and dune-capped barriers with lagoons. Directly resting on the crystalline rocks, sedimentary deposits largely constitute the coastal geomorphology around Tolanaro. Three successive transgressions during the Quaternary formed the landscape of the coast, leaving areas of sand dunes and calcarenites. On cliffy sectors shore platforms show solution basins (plateformes à vasques). Annual rainfall increases from 300–700 mm on the south-southwest coast to more than 1,500 mm in the southeast, but mountain slopes in the lee of the wind are relatively dry. The east coast of Madagascar, between Tolanaro and Toamasina, is low, sandy, and exceptionally straight for about 800 km; it is somewhat inhospitable. Its general outline corresponds to one of the major tectonic lineaments of Madagascar, running north-northeast to southsouthwest, but it is not a fault scarp. The continental shelf is narrow, with a slope descending into oceanic depths of

more than 3,000 m. In detail, the coastline has been smoothed by the action of ocean swell, which has built along its length a sandy Holocene outer barrier consisting of numerous beach berms (>Fig. 18.1.9). Coral reefs are limited (occurring at Manakara and off the port of Tamatave), and mangroves are rare. It is a lagoon coast, the chain of lagoons lying between the outer barrier and an earlier Pleistocene (Eemian) barrier con­sisting of white, leached sand, 8–12 m above present sea level. The climate is ­characterised by high precipi­tation and temperatures. Furthermore, the trade winds bring high atmospheric humidity and rainfall from the Indian Ocean. From January to April tropical cyclones are common. The Canal des Pangalanes has been cut through lowlying segments between the lagoons to form a narrow waterway that supports traditional inland navigation, but is otherwise little used. Between Manakara and Mananjara (>Fig. 18.1.10) long stretches of beach, brackish lagoons, and wide sand ridges are typical of the coastline. Built on sandy beach ridges, Toamasina (formerly Tamatave), the main port of Madagascar, is partly protected by coral banks with small cays. Marine erosion is a major problem for this provincial capital.

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⊡⊡Fig. 18.1.9 The sandy east coast north of Toamasina, with plantations of Casuarina equisetifolia on beach ridges. (Courtesy P. Le Bourdiec.)

North of Toamasina the straight coast ends at Foulpointe, beyond Mahavelona, and a bay extends past Fenoarivo. A sandy coast continues to Pointe a Larree, a  cuspate foreland in the lee of Ile Sainte-Marie (Nosy  Borah), an elongated island, mainly of granite. Further north the coast becomes hilly on granite, with outlines determined by two directions of faulting: northnortheast to south-southwest and north-northwest to south-­southeast. Antongila Bay occupies a graben between faults running in the latter direction. On its eastern side granitic slopes rise from a rocky shore that extends round the Masoala Peninsula and up to Cape Est. The coast is then low and sandy past Sambava. Fringing coral reefs are extensive at Mahambo, Fenoarivo, and Mananara, as well as on Ile Sainte-Marie. The low, sandy coast continues past Vohimarina, and steepens on the Amalamera Jurassic limestone north of Baie de Loky. There are inlets at valley mouths invaded by the Flandrian marine transgression. The Rade of Diégo-Suarez (now Antsiranana), close to the old military port, is a magnificent fringing ria with a narrow entrance. It is one of the best-sheltered naval anchorages in the world. ⊡⊡Fig. 18.1.10 Mananjara: river mouth and sand spit.

References Battistini R (1995) The Sarodrano spit (in French). Norois 42, 165:63–71 Lebigre JM (1997) L The mangroves of the southwest of Madagascar (in French). In: Lebigre JM (ed) Milieux et sociétés dans le Sud-Ouest de Madagascar, vol 23. CRET, “Iles et Archipels”, pp 135–242

18.2  Indian Ocean Islands

Wong Poh Poh

1.  Introduction Some Indian Ocean islands are described in other chapters: > Madagascar > Sri Lanka and > Indonesia while subantarctic islands south of 50°S are described in > Antarctic Coast. The islands of the Indian Ocean (>Fig.  18.2.1) fall into three major groups: the numerous archipelagoes of the western part, mainly north and east of Madagascar; the chain of reef islands on the submarine Maldives Ridge, west and southwest of India, extending from the Laccadive and Maldive islands south to Diego Garcia; and the northeastern group, including the Andaman and Nicobar Islands, the Kepulauan Mentawai west of Sumatra, and the outlying Christmas Island and the Cocos Islands. There are three kinds of islands: high islands (either

v­ olcanic or granitic); islands of raised coral reefs bordered by low cliffs; and low cays of coral sand, usually with beach rock of recent origin. Cays are produced by Holocene deposition of coral sand and gravel on an intertidal coral reef platform. They rise above normal high tide level, and may be colonised by vegetation. Their shores consist of wide beaches, sometimes with storm-placed coral blocks and outcrops of recent beach rock (Stoddart 1970; Stoddart and Yonge 1971).

2.  The Comores Archipelago The Comores archipelago, in the north of the Mozambique Channel and northwest of Madagascar (12°S, 44°E), includes four high islands: Grand Comore, Moheli,

⊡⊡Fig. 18.2.1 Islands of the Western Indian Ocean. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_18.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡Fig. 18.2.2 Grand Comore, in the Comores Archipelago.

Anjouan, and Mayotte. Grand Comore (>Fig. 18.2.2) has an area of 1,131 sq. km, and consists of a large active volcano, Kartala (2,361 m), to the north of which is attached a chain of Strombolian cones of Upper Quaternary age, forming the Massif de la Grille peninsula. The coastline comprises low cliffs (>Fig.  18.2.3) cut into Holocene basaltic lava flows, with occasional sandy beaches generally associated with fringing reefs in the north (Mitsamiouli), in the west (Moroni and Itsandra), in the east (>Fig.  18.2.4), and in the south (Male and Chindini). Sand mining has been a common practice leading to much erosion and severe loss of beaches (>Fig. 18.2.5). Sea­walls have been built to protect some local communities, as at Bangoua Kouni, Ntsaoueni. Mangroves are still present on the more sheltered south coast; elsewhere, some have managed to establish on thin sediment on low rock platforms (>Fig. 18.2.6). Anjouan, to the southeast, and Moheli (>Fig. 18.2.7) are older (Pliocene) volcanic islands, much dissected by erosion, and rising to peaks of 1,595 and 790 m, respectively. Their coastlines are more eroded, with sectors of cliff, some­times high, cut in lavas, alternating with sandy bays that include small mangrove areas. On Moheli the protected beach at Itsamia is one of the few important nesting areas for sea turtles in the Indian Ocean; columnar basalts are also exposed (>Fig. 18.2.8). Coastal erosion has led to the construction of seawalls to protect local communities, as at Nioumachoua and Domoni. Mangroves have spread as the result of accelerated accretion from rivers bringing down sediment from ⊡⊡Fig. 18.2.3 Aerial view of rocky coastline immediately south of the Grand Comore airport.

Indian Ocean Islands

18.2

⊡⊡Fig. 18.2.4 Calcareous sandy beaches between rocky headlands, Bouni.

⊡⊡Fig. 18.2.5 Severe erosion arising from beach mining, Ndroude.

hinterlands where soil erosion is widespread, such as in  the bays east and west of Nioumachoua. Moheli is ­virtually ringed by narrow fringing reefs, and there were reports of severe coral bleaching during the 1983 El Niño event. Mayotte (>Fig.  18.2.9) is a much dissected volcano rising to 660 m, partly of Miocene age, with deep embayments such as Boueni, and a surrounding lagoon up to 15 km wide enclosed by a 140 km barrier reef, which in

the south becomes a double barrier (Guilcher et al. 1965). None of these islands shows Pleistocene raised beaches. The island of Lys (Glorieuses) in the Mozambique Channel is an emerged coral reef (Battistini and Cremers 1972), and there are a number of sand cays on reefs in the Mozambique Channel, notably Juan de Nova, west of Madagascar, and Europa and Bassas da India, which are small Eemian atolls, with shallow lagoons occupied by mangroves (Battistini 1966).

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⊡⊡Fig. 18.2.6 Young mangroves on a rock platform, Domoni, Grand Comore.

⊡⊡Fig. 18.2.7 Moheli, in the Comores Archipelago.

3.  The Mascarene Archipelago The Mascarene archipelago is formed by three high volcanic islands – Réunion, Mauritius, and Rodriguez – along the sea floor ridge between Mauritius and the Seychelles. Réunion (>Fig. 18.2.10) is a large Hawaiian-type volcano that includes an older part, the massif of Piton des Neiges (3,069 m), incised by three large cirques, and an active volcano to the southeast, the massif of Fournaise (CazesDuvat and Paskoff 2004). Much of the Réunion coast is rocky, with low cliffs cut in lava. The cliffs are higher between Saint-Denis and La  Possession on the north. The rocky coastline varies

from typical cliffs (>Fig.  18.2.11) to low headlands (>Fig. 18.2.12). Sectors of low coast correspond with large depositional cones below three cirques – Malfate, Salazie, and Cilaos – where gravel beaches are found (>Fig. 18.2.13). The limited sandy beaches are related to embryonic fringing reefs on the west coast: Boucan Canot where some erosion has occurred, Saint-Gilles where its southern limit is marked clearly by a natural rock groyne at Pointe Trois Roches, Saint-Leu where some beach rock is exposed (>Fig. 18.2.14), Etang Sale with its cuspatespit, and SaintPierre. Sand winnowed from beaches produces dunes at Baie de Saint-Paul and at Etang du Gol, where they reach a height of 4–5 m.

Indian Ocean Islands

18.2

⊡⊡Fig. 18.2.8 Columnal basalts on the rock platform at Itsamia.

⊡⊡Fig. 18.2.9 Mayotte, in the Comores Archipelago. PRE - MIDDLE PLEISTOCENE VOLCANIC ROCKS (RECENT VOLCANICS BLANK) GRAVEL CONES CORAL REEF SANDY COASTLINE

M ZAMBOUROU

ROCKY COASTLINE M MANGROVE

M

M LAGOON M B. DE BOUENI

PAMANZI M

572m

660m

M

M M M

N

MAYOTTE

0

5

10 km

4.  Mauritius Mauritius (20°S, 58°E), 180 km northeast of Réunion, is a less elevated island (826 m at Piton de la Rivière Noire) with

more varied relief (>Fig. 18.2.15). It has an area of 1,860 sq. km, and consists of a much-dissected Pliocene volcano, with residual escarpments from which has emerged a newer, flatter Quaternary volcano composed of basaltic flows of faint relief. The coast is generally low, cut in the basalts. The highest cliffs are in the south (>Fig. 18.2.16), with low headlands elsewhere, a few with distinctive patterns of polygonal columnar joints (>Fig. 18.2.17). The entire coastline of Mauritius is reef-rimmed up to a maximum distance of 4 km from the shore. This has resulted in a number of major sandy beaches on the north coast (Cap Malheureux to Trou aux Biches), the east coast (Pointe de Flacq to Pointe Quatre Cocos), and the west coast (Flic en Flac to Tamarin). A notable feature is the absence of reef flats, thus explaining sediment accumulation in lagoons (Wafar et al. 2003). This also partially explains the formation of muddy coasts in sheltered large bays in the northeast, Mahebourg, and north of Le Morne Brabant. Such areas are suitable for barachois or bodies of water partially enclosed by artificial embankments and sluice openings to trap fish brought in by the tides. Distinct river mouth barriers have formed across drowned valleys on the west coast (>Fig.  18.2.18). Dunes are found in a number of locations, and aeolianite at several locations in the south. Rapid erosion of beaches occurs during tropical cyclones, as in 1960, when many beaches were scoured away or swept landward (McIntire and Walker 1964).

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Indian Ocean Islands

⊡⊡Fig. 18.2.10 Réunion, in the Mascarene archipelago.

⊡⊡Fig. 18.2.11 Cliff at Cap La Houssaye.

There are also the remains of old emerged coral reefs, some slightly recalcified at 3 m, other strongly recalcified at 10 m, indicating recent uplift, at least in the southern part of the island.

Rodriguez, further to the northeast, is also a volcanic island, bordered on the south, west, and north by fringing reefs 3–8 km wide, enclosing a shallow lagoon (less than 2 m deep at low tide).

Indian Ocean Islands

18.2

⊡⊡Fig. 18.2.12 Low rocky promontory at Pointe Finition.

⊡⊡Fig. 18.2.13 Gravel beach ridge at Etang de Saint-Paul.

5.  Seychelles The Republic of Seychelles includes 42 granitic islands  and 74 coralline islands, with a total area of 455  sq km. The climate is humid tropical: Victoria, the  capital, has mean monthly temperatures of 26.9°C in  January and 25.8°C in July, and an average annual

rainfall of 2,375 mm. In ­general, the granitic islands are higher and wetter, and carry a richer vegetation than the coralline islands. The most important of the granitic islands is Mahe, 30 km long and up to 906 m high, with a steep eroded coastline where rocky sectors alternate with sandy beaches. Rocky sectors vary from distinct granite outcrops with

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Indian Ocean Islands

⊡⊡Fig. 18.2.14 Beach rock exposed at Saint-Leu, Réunion.

⊡⊡Fig. 18.2.15 Mauritius, in the Mascarene archipelago.

Indian Ocean Islands

18.2

⊡⊡Fig. 18.2.16 The Souffler, a blowhole on the cliffy south coast of Mauritius.

⊡⊡Fig. 18.2.17 Polygonal basalt truncated on a rocky shore at Cape Malheureux, Mauritius.

some pseudokarren features (>Fig. 18.2.19) to low headlands separating the bays. A fringing reef with an average width of 600–800 m has developed on the eastern seaboard of Mahe (Wagle and Hashimi 1990). Straight beaches are rare, and crescentshaped beaches vary from small to large bays. Erosion of beaches is common, and more intense on the western side

of Mahe. This explains why beach rock is more widespread on the east coast than on the west coast, as coastal retreat is slow and beach rock is preserved (Wagle and Hashimi 1990). Mangroves are limited to several patches, mainly on the west coast. A rare example of mangrove regrowth takes place in the backwater areas around land reclaimed between Victoria and the international airport (Taylor et al. 2003).

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Indian Ocean Islands

⊡⊡Fig. 18.2.18 Sandy barrier across the mouth of Rivière Tamarin, Mauritius.

⊡⊡Fig. 18.2.19 Prominent granite outcrop at Cap Lascars.

The coralline islands and atolls lie south and west of the granitic islands. The four islands of the Aldabra group in the southern Seychelles are slightly elevated reefs, of Pleistocene (Eemian) age at Aldabra. There are cliffs up to 8 m high. At Aldabra (>Fig. 18.2.20) the Eemian atoll is well preserved, the Holocene marine transgression having merely reoccupied the older lagoon,

and the surrounding coastline is cliffed, cut in older coral limestone (Stoddart 1967). Cosmoledo is an old emerged reef in the form of scarped rocky islets. Astove is a small raised annular reef, while Assumption is an old cay-capped platform, uplifted and fringed by rocky reef cliffs up to 7 m high. On the south and east coasts are dunes up to 30 m high. In the

Indian Ocean Islands

18.2

⊡⊡Fig. 18.2.20 Albadra. (Courtesy Geostudies.)

southern Seychelles, Saint-Pierre Island is a circular emerged coral atoll with cliffs up to 10 m high. The Amirantes in the western Seychelles are small Holocene sand cays on atolls and reef patches, and there are sandy islands on the eastern reef of the atoll of Farquhar (Battistini and Jouannie 1979). Socotra (8°N, 53°E) is a large (3,582 sq. km) island, with a mountainous interior rising to 1,520 m. The coastline partly consists of steep limestone cliffs rising 500– 600  m to an undulating plateau, separated by low-lying coastal plain sectors.

6.  Laccadive and Maldive Islands The Laccadive and Maldive islands stand on the submerged Maldives Ridge, which runs from north to south off the west coast of India. They include hundreds of small cays arranged in chains, and atolls of larger dimensions. The Lakshadweep Islands are a group of 12 atolls and 3 reefs bearing a total of 36 cays. The Maldives consist of more than 1,200 islands in 19 groups of atolls, and are entirely low-lying (>Fig.  18.2.21). Malé, the capital city, occupies a reef island on the rim of a large atoll. Reef rock has been quarried from the bordering reef for the construction of an enclosing sea wall, so that the high tide shoreline is entirely artificial around an urbanised island 1.25 km long and 800 m wide, standing a metre above sea level. Outside the sea wall the ocean floor declines steeply to deep water. Construction of another artificial walled island, Hulhumale, began in 1997. Linked by a causeway to the international airport, this island is being built 2 m above sea level. To the south Addu and the low islets of the Chagos Bank are sand cays on reef platforms, mainly on five

atolls (> Fig. 18.2.22). Some of the lithified coralline sands may be of pre-Holocene origin. Two atolls, southern Peros Bahos and the northwestern part of Great Chagos Bank, include emerged reefs bordered by cliffs up to 6 m high in reef limestone. Diego Garcia is an atoll with four low islands at the southern end of the island chain. In the eastern Indian Ocean, about 960 km southwest of Sumatra, the Cocos (Keeling) Islands (12°S, 97°E) are a chain of 27 coralline islands on atoll reefs encircling a lagoon. There are heavy ocean swells and frequent storms. Christmas Island, to the east (10°35ʹ S, 105°; 35ʹ E), is a strongly uplifted coral reef, culminating at 350 m, and bordered by high limestone cliffs. Its extensive guano deposits have been heavily exploited for phosphates.

7.  Andaman and Nicobar Islands In the northeast Indian Ocean (Andaman Sea) the Andaman and Nicobar Islands are a chain of high, mainly forested islands, mostly of sedimentary rock, essentially the unsubmerged peaks of a southward continuation of the Arakan Yoma Range in western Burma. Deep valley-mouth inlets, submerged during the Holocene marine transgression, remain unfilled because of meagre sediment yields from small river catchments. Steep coasts are typical, cliffed on the more exposed westerly shores, with beaches, spits, and mangrovefringed bays and inlets on more sheltered eastern sectors. Fringing reefs are extensive. Emerged beaches and shore platforms, and uplifted reefs have been noted at various elevations of up to 5 m, with evidence of tectonic deformation, while to the east Narcondam (710 m) and Barren Island (353 m) are of volcanic origin.

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Indian Ocean Islands

⊡⊡ Fig. 18.2.21 The Maldive Islands, with detail of Fadiffolu. (Courtesy Geostudies.)

Between the Andaman and Nicobar Islands is the Ten Degree Channel, noted for its strong currents. The Boxing Day 2004 Indian Ocean tsunami had a major impact here.

8. Impact of the 2004 Indian Ocean Tsunami The 26 December 2004 Indian Ocean tsunami, caused by an earthquake of M9 off the west coast of Sumatra, was a rare event that affected the coasts of the Indian Ocean islands. As they are in the same subduction zone of the earthquake, the Nicobar and Andaman Islands were severely affected by both the earthquake and the tsunami. Subsidence occurred on the eastern coast of South Andaman Island and all the Nicobar Islands, while uplift occurred in the North and Middle Andaman islands. Rock platforms and coral beds were exposed on the western side as a result of relative uplift. The Maldives are well protected by coral reefs and the tsunami did not cause any substantial erosion, but accentuated predictable seasonal oscillations in shoreline changes (Kench et al. 2006). However, the tsunami waves severely affected the coasts of granitic islands in the Seychelles farther west, resulting in erosion and damage to property and infrastructure, in some cases due to the removal or lowering of the beach berms. Farther to the south, the other Indian Ocean islands, Mauritius, Reunion, and the Comores, were much less impacted by the tsunami. ⊡⊡Fig. 18.2.22 A cay in the Maldives. (Courtesy Ken Boston.)

Indian Ocean Islands

References Battistini R (1966) La morphologie de l’ile Europa. Musée Nationale d’Histoire Naturel, Memoir A-41:7–18 Battistini R, Cremers G (1972) Geomorphology and vegetation of Iles Glorieuses. Atoll Res Bull 159:1–10 Cazes-Duvat V, Paskoff R (2004) Les littoraux des Mascareignes entre nature et aménagement. Harmattan, Paris Guilcher A, Berthois L, Le Calvez Y, Battistini R, Crosnier A (1965) Les recifs coralliens et le lagon de 1'ilε Mayotte (Archipel des Comores, Ocean Indien). ORSTOM, Paris Kench PS, McLean RF, Brander RW, Nichol SL, Smithers SG, Ford MR, Parnell KE, Aslam M (2006) Geological effects of tsunami on midocean atoll islands: the Maldives before and after the Sumatran tsunami. Geology 34:177–180

18.2

McIntire WG, Walker HJ (1964) Tropical cyclones and coastal geomorphology in Mauritius. Ann Assoc Am Geogr 54:582–596 Stoddart DR (1970) Coral Islands of the Western Indian Ocean. Atoll Res Bull 136:1–224 Stoddart DR, Yonge M (1971) Regional variation in Indian Ocean coral Reefs. Academic, London Taylor M, Ravilious C, Green EP (2003) Mangroves of East Africa. UNEP 24 pp Wafar M, Wafar S, Yennavar P (2003) Indian Ocean Islands, coastal ecology and geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Kluwer, Dordrecht, pp 557–564 Wagle BG, Hashimi NH (1990) Technical note – Coastal geomorphology of Mahe Island, Seychelles. Int J Remote Sens 11:281–287

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19.0  South-East Asia – Editorial Introduction

The mountain ranges of Palaeozoic rock that run along the border between > Myanmar and Thailand continue through the isthmus to the south, one range running out to Phuket Island. There are intrusions of late Cretaceousearly Tertiary granite, exposed on the west coast, in the Samui Islands, at Ko Kra, and at Narathiwat. Outcrops of Permian limestone form the remarkable karstic islands in Phang Nga Bay, and occur at intervals southward into Malaysia. The Andaman Sea coast of Thailand is generally steep and locally cliffed, with sandy beaches and spits, and mangroves and tidal mudflats in sheltered bays and inlets. The coast receives swell from the Indian Ocean, and waves generated by southwest monsoonal winds, and the tide range is about 2 m. Peninsular > Malaysia is threaded by ranges produced by late Jurassic folding, consisting of Devonian formations and granites on the west coast and Permian and Carboniferous rocks to the east. The west coast is sheltered by Sumatra, and Indian Ocean swell fades out south of the rocky island of Langkawi. Mangrove swamps are extensive, particularly around deltas such as that of the Klang River, but there are headlands where ridges of Devonian rock come to the coast. Mean spring tide ranges are generally about 2 m, increasing to 4.1 m on the Klang delta. Low wave energy permits mangrove swamps on the west coast of the granitic island of Penang. > Singapore is a hilly island of granite and Mesozoic sedimentary rocks, with coastal lowlands and a coastline largely artificial as a result of land reclamation. The east coast has wide coastal plains with several deltas, notably those of the Pahang, Terengganu, and Kelantan rivers, and sandy beaches often backed by parallel beach ridges (permatang). Wave action from the South China Sea is strong during the northwest monsoon. Mean spring tide ranges are generally between 1.5 and 2 m, diminishing northward. The coast of the Gulf of Thailand is generally low-lying and fringed by sandy beaches, with mangroves in inlets and estuaries and headlands with steep slopes and cliffs at the terminations of hilly ridges. At the head of the Gulf the Bight of Bangkok coast is a deltaic lowland with ricefields behind a narrow sandy beach, and the mouth of the Chao

Phraya River. To the south-west the coast becomes steeper, with cliffed promontories between bays backed by sandy beaches and beach ridges, and mangroves in inlets. Similar features are seen on the coasts of > Cambodia. The Mekong Delta dominates southern > Vietnam, but the east coast has a generally narrow coastal plain (up to 30 km wide) backed by a mountainous hinterland. Numerous rivers flow down from the mountains and across the narrow coastal plain into estuaries that interrupt long sandy beaches. Cliffed headlands mark the terminations of hilly ridges that cross the coastal plain, and there are occasional outcrops of Permian limestone. The > Philippines is a mountainous archipelago, with many sectors of steep coast, locally cliffed, and beaches in bays. It has been disturbed by neotectonic movements and active volcanoes. Mangroves grow in sheltered situations, mainly in estuaries and bays, and coral reefs are widespread. The tide range is generally less than 2 m. The eastern coast is exposed to Pacific Ocean swell, as well as to  storm waves generated by northeast monsoon winds and occasional typhoons. The western coast is exposed to smaller waves from the South China Sea, but also to typhoon surges. Within the archipelago there are many low wave energy sectors on coasts facing only short fetches. The > Indonesian archipelago includes Sumatra, which has a steep and mountainous southwestern coast with a chain of hilly islands offshore, and a wide, swampy eastern lowland with mangroves fringing estuaries and inlets. Volcanic formations (with some active volcanoes) are a major formation of the ranges that continue through Java, where they are bordered by northern deltaic plains, to Bali and the chain of high islands to the east, including > East Timor. Their southern coasts have extensive coral reefs and cliffs cut in emerged coral limestone. Beaches receive swell from the Indian Ocean, whereas the northern coasts are sheltered, subject to limited wave action produced by monsoonal winds on the Java and Flores seas. The archipelago is tectonically active along a plate margin, and terraces at several levels result from intermittent uplift. Tide ranges are less than 2 m, and often below 1 m.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The island of Borneo (including the Malaysian Sabah and Sarawak, > Brunei, and Indonesian Kalimantan) has a high hinterland, with some ridges running down to coastal headlands, but much of the coast is low-lying and beachfringed, backed by beach ridges (permatang). Several rivers drain to the swampy south and east coasts, but form steep coasts in the north mountain ranges in Sabah. Much of the intricate coast of Sulawesi is steep and mountainous. The northern coasts of Irian Jaya and Papua New Guinea have steep, mountainous sectors, exposed to Pacific Ocean swell and waves generated by local monsoon winds. There are bordering beaches of sand and gravel, supplied

mainly from rivers draining steep ­catchments, and coral reefs are extensive. Marine terraces indicate Quaternary tectonic uplift, as on the Huon Peninsula, which has a stairway of emerged coral reefs. The tide range is small (less than 1 m). The south coast of Irian Jaya is low-lying, with extensive swamps, and to the east in Papua the Fly River delta has mangrove-fringed estuaries and islands. The coast is steeper to the east, past Port Moresby, on the southern slopes of Owen Stanley Range, where ridges run out to headlands between bay beaches. Fringing coral reefs occur along this coast, which to the east is protected by the barrier reef that runs on around the Louisiade Archipelago.

19.1 Thailand Andaman Sea Coast

Sanit Aksornkoae · Eric Bird

1.  Introduction The coastline of Thailand is about 2,960 km long, of which 750 km is on the west (Andaman Sea) coast, 1,670 km on the Gulf of Thailand, and the remainder (520 km) on over 258 islands, of which the six largest account for 90% (Pitman 1982). >Table 19.1.1 summarises the distribution of major coastal types. In order to maintain the roundthe-world sequence the Andaman Sea coast is presented here, and the Gulf of Thailand coast after Malaysia and Singapore (> Gulf of Thailand). The outline of central and peninsular Thailand is controlled by the effects of an Upper Palaeozoic Hercynian orogeny, modified by a Mesozoic orogeny, which produced fold belts that trend north in the south but northwest toward the head of the Gulf of Thailand. These Cambrian to Triassic sediment were intruded by late Cre­ taceous to early Tertiary granitic masses, which form four major linear ridges: the first, adjacent to the west coast forms Phuket Island and extends to Ranong and Kan­ chanaburi; the second, the Satun ridge, gives rise to the islands of the Samui group; while the third and fourth form prominent ridges and headlands at Ko Kra and Nara­ thiwat. Northeast of Phuket Island, the peninsula is crossed by a complex of northeast–southwest faults, interpreted as major transcurrent faults that displace the Phuket granites sinistrally 250 km to the southwest along the Khlong Maru

fault. The granitic outcrop in southeast Thailand is similarly displaced. Between these ridges, particularly the Phuket and Samui massifs, major faulting has produced a graben, in which Permian Ratburi Limestones are found, responsible for unique marine karstic landforms in Phang Nga Bay, down to the Malaysian border and along the Gulf coast as far as north Petchaburi. Tertiary sediments are confined to small basins and Tertiary orogenic movements resulted in the subsidence of the Chao Phraya depression. The major directions of alongshore drift reflect the  interaction between winds, waves and currents. The Andaman Sea coast receives ocean swell from the Indian Ocean. The southwest monsoon, with winds mainly from the west and south is from May to October. Maximum annual rainfall is over 2,200 mm on the west coast. Flooding is common. Mean spring tide ranges on the west coast are between 1.9 and 2.8 m. Tidal mudflats, inundated daily by the sea, are extensive in Phang Nga Bay. Extensive and luxuriant growth of mangrove forest has occurred along the Andaman Sea coast. In all, there are some 1,620 sq km, 875 sq km on the Andaman Sea coast (Watson 1928), where extensive tracts are found behind the broad tidal mudflats, particularly between Satun and Phang Nga and north of Takua Pa to Ranong. On the Andaman Sea coast individual mangrove trees may be of 40 m high. These swamps are marked by clearly defined

⊡  Table 19.1.1 Distribution of major coastal types in Thailand (km). (Courtesy Geostudies.) Area Type

Andaman Sea

Eastern Peninsula

Mangrove/MudFlat

480.0 (55.7%)

103.8 (12.0)

Sand Beach

207.5 (16.6%)

Rocky Coast

106.0 (24.1%)

Total

793.5

Head of Gulf

East Coast

Total

168.0 (19.5%)

110.0 (12.8%)

862.0 (33.0)

795.0 (63.4%)

64.0 (5.10)

187.5 (14.9%)

1254.0 (49.0)

241.2 (54.7%)

7.0 (16%)

86.5 (19.6%)

1140.0

239.0

384.0

440.7 (17.2%) 2556.7 (100)

A revised version of a chapter by J.I. Pitman in The World’s Coastline (1985: 771–787). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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ecological zonations, forming five major community types, related to frequency and depth of flooding (Aksornkoae 1976). Seaward of the mangroves are extensive mudflats, with gradients of less than 1%. Landward these estuarine deposits merge into extensive alluvial deposits of about 5 m above sea level. Mangroves have been cleared for timber, fuel and charcoal, while nipa is used for roofing. Current estimates are that 80% of the mangrove areas in Thailand have been converted to prawn mariculture. Sandy beaches are found mainly where there is a supply of sand from rivers or nearby cliff erosion, and also where sand has been swept in from the sea floor, as on the Phuket coast. Coralline sand and gravel beaches are found near eroding reefs. Coral reefs occur on almost all the islands off the Andaman Sea coast, other than those with mangroves or muddy shores, and are being developed for tourism around Phuket Island. Ko Samui and Ko Phangna have extensive reefs, especially on their southern and eastern coasts, while the smaller islands of this group have surrounding reefs.

2.  Andaman Sea Coast South from the Burmese borders the Thai coast consist of many mangrove-fringed bays and inlets between peninsulas that end in a few bluffs or cliffs, mainly of weathered rock. Ranong is an old port town on an estuary opening to the strait behind Victoria Point in Burma. To the south,

steep forested granite hills run behind a sandy pediment that descends to the coast. Prawn ponds have been excavated on land just behind the mangrove fringe. The mangroves become less extensive, as the coast becomes more exposed to waves from the Andaman Sea, and high forested hills form steep capes between bays lined by beaches of fine grey sand, backed by faint beach ridges under Casuarina woodland (>Fig. 19.1.1). Low ocean waves with periods of 12–16 s move in to the wide, sandy prograding Pra Prat Beach. Northward drifting has supplied sand to Lem Tom Chob spit, and tin is excavated from beach ridge deposits on the coastal plain. A forested range on dark slates, sandstones and schists runs south to Phuket Island, which has a west coast with sandy surf beaches between rocky promontories. There is a monsoon wave trimline on headlands and islands, notably in Nai Harn Island. Coral reefs develop towards the southern end, and beach roch outcrops locally after storm wave erosion. Tsunamis are rare, but a major one occurred on Boxing Day 2004, as a result of an earthquake south of Sumatra. After an initial withdrawal, a series of waves up to 5 m high broke on the shores of Phuket and the Andaman Sea coast, causing extensive sea flooding and structural damage. On several sectors, sand was washed onshore by these waves and deposited on and behind the pre-existing beaches. On the east coast of Phuket Island, limestone ridges up to 300 m high run out as rugged islands and stacks in Phangnga Bay (>Fig. 19.1.2), a fine example of tower karst

⊡⊡ Fig. 19.1.1 Sandy beach and low beach ridges under Casuarina woodland on the Andaman Sea coast south of Rayong. (Courtesy Geostudies.)

Thailand Andaman Sea Coast 

topography. Vertical limestone cliffs are partly vegetated down to a basal notch (>Fig. 19.1.3), and in places, there are remainders of an older notch cut when the sea stood about 3 m above its present level, with dripstones from the overhanging visor and the roofs of caves. There are pale scars on cliffs where rock masses have fallen away, and pocket beaches in coves. One island has a tombolo beside a large tilted slab of limestone. At the head of Phangna Bay, are mangrove-fringed creeks. The Phi Phi islands are also very steep sided limestone islands with basal notches at sea level. There are occasional sandy beaches, some with low dunes. One of these, in Maya Bay on Phi Phi Leh, was modified in 1999 by bulldozing the dunes to make a wide sand plain for the making of the film titled ‘The Beach.’ In the next monsoonal storm the sand was washed away, but the Indian Ocean tsunami in December 2004 restored the beach, and the dunes re-formed. To the east is the Than Phut lowland with limestone ridges which end in peninsulas, stacks and outlying islands such as Ko Yao Noi and Ko Yao Yai. Some islands are tied to the mainland by sandy tombolos, others have lee spits paired with spits protruding from the mainland. At Shelly Beach, layers of shelly Miocene limestone and lignite dip into the sea as structural ledges. As a result of recent submergence, the steep offshore gradient and the effects of the southwest monsoon, coastal sediment are restricted to river mouth estuaries, behind bay-mouth barriers, and in sheltered locations eastward of offshore islands. Mangrove swamps extend up to 10 km inland along tidal creeks and ⊡⊡ Fig. 19.1.3 Basal notch on cliff in Phangnga Bay. (Courtesy Geostudies.)

   19.1

⊡⊡ Fig. 19.1.2 Limestone island in Phangnga Bay. (Courtesy Geostudies.)

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estuaries, gradually giving way to fresh-water vegetation. These swamps are marked by anastomising channel patterns, flanked by steep mudbanks. The mangroves have been cut for fuel and dredged for tin. Generally the peaty swamp deposits are 3–5 m deep, overlying gravels and alluvium. The indented, island-fringed coast continues south east to Satun and the Malaysian border, with extensive mangroves, especially in inlets and estuaries.

References Aksornkoae S (1976) Structure, regeneration and productivity of mangroves in Thailand. PhD thesis, Michigan State University, East Lansing, MI Pitman JI (1982) Geomorphology of the coastline of Thailand. Occasional paper, Department of Geography, King’s College, London Watson JG (1928) Mangrove forests of the Malayan Peninsula. Fraser and Neave, Singapore

19.2  Malaysia – Introduction

Teh Tiong Sa · Yap Hui Boon   

1.  Introduction Malaysia is a maritime nation, made up of the 13 states of Peninsular Malaysia (sometimes referred to as West Malaysia), Sabah and Sarawak (sometimes referred to as East Malaysia), and the Federal Territories of land-locked Kuala Lumpur and the island of Labuan (>Fig.  19.2.1). According to EPU (1985), the Malaysian coastline (not counting the smaller islands) is 4,809 km long, made up of 1,972 km in Peninsular Malaysia, 1,035 km in Sarawak, and 1,802 km in Sabah (including Labuan). In the peninsula the more indented west coast facing the Straits of Melaka is longer than the straighter east coast facing the South China Sea. More than 90% of the coastline is made up of easily erodible material, sand or silt and clay. Sandy beaches are most prevalent on the coasts facing the South China Sea, while mangroves are most prevalent along the Straits of Melaka and on the coasts facing the Sulu and Celebes Seas. 51% of the coastline is sandy, 42% has mangroves, 6% is rocky, and 1% is man-made. The proportion of man-made shorelines is increasing at the expense of mangroves, largely due to coastal land reclamation and the construction of coastal structures to combat erosion. Beaches are generally sandy, but gravel beaches are found where pebbles and cobbles have been derived from intricately fissured rocky outcrops, conglomerates, or gravels in coastal terraces. Sandy beaches have been supplied with sand carried down to the coast by rivers or swept in from the sea floor, but on the coast of Penang there are sandy beaches formed of material derived from the weathered mantle of the island granites. On the peninsula mangroves protect the entire length of the sheltered west coast, but are confined to estuaries and behind spits on the east coast. According to Teh and Lim (1993), slightly more than half of the mangrove coasts in the peninsula are retreating and 11.2% of these are receding at more than 8.0 m/year, all of which are recorded in Penang, Perak, and Selangor. In Sabah, mangroves are the dominant coastal landform, and are found extensively in Marudu Bay and along the coast facing the Sulu Sea. In Sarawak large areas of mangroves are found in Datu Bay. Similar to the rest of Southeast Asia, large areas of

mangroves in Malaysia had been converted to agriculture (Bird and Teh 2006). Some reclaimed mangrove areas have been abandoned (Teh and Bird 1999). Actively receding cliffs are rare in Malaysia, but occasionally there is slumping of the weathered mantle, especially after the foot slope has been undercut by storm waves. However, active cliffing is seen on the more exposed shores of promontories and islands, washed by waves generated by the monsoons. Headlands on coastal outcrops of granite and sedimentary rocks usually have steep slopes mantled with weathered material, held in place by a scrub and forest cover. In the case of granite slopes, there is often an apron of core boulders protecting the base. On the peninsula, cliffs are found mainly on offshore islands such as on Tioman, Penang, and Pangkor, where granite cliffs form steep slopes characteristically bordered by aprons of core stone boulders. The plunging cliffs of Langkawi are cut in limestone or quartzite. In Sarawak, steep sandstone cliffs have developed on the Bako Peninsula, and on Berhala Island at the entrance to Sandakan Bay. The cliffs south of Miri are cut into elevated terraces. Man-made coastlines are becoming increasingly common. In fact, if bunded coastlines are included in this category, they now form a major coastal type on the west coast of Peninsular Malaysia.

2.  Coastal Setting Sea levels along the coast of Malaysia are influenced mainly by the co-oscillating tides of the Pacific and Indian basins. Thus, the region has considerable variation in tidal types. In Peninsular Malaysia the tides are mixed on the east coast and semi-diurnal on the west, whereas tides in Sabah and Sarawak are usually mixed, except for diurnal tides between Bintulu and Kuala Baram. In general, mean spring tidal ranges around the peninsula are small, up to 2.7 m in the Straits of Melaka and less than 2 m on the east coast, the exception being in the region of the Kelang delta (4.1 m in Port Kelang). Along the east coast Kuantan records 1.9 m, with the range diminishing northward to 0.8 m in Tumpat. In Sarawak and Sabah tide ranges are generally less than 2 m, but attain 3.6 m on Pulau Lakei in

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Malaysia – Introduction

⊡⊡ Fig. 19.2.1 Malaysia, comprising West (Peninsular) Malaysia and East Malaysia Sarawak and Sabah.

Sarawak and 2.5 m at Tawau on the east coast of Sabah. The highest astronomical tide recorded on the east coast of Peninsular Malaysia is at Tg. Gelang (1.8 m above mean sea level), and the highest in Sabah and Sarawak at Kuching (2.7 m above mean sea level). Storm surges and wave set-up can significantly affect sea surface elevation along the Malaysian coast, especially along the coasts facing the South China Sea. The storm surge for a 30-year recurrence period varies from a low 0.39 m in Morib, on the sheltered west coast of Peninsular Malaysia, to 1.36 m in Miri, facing the South China Sea. In addition, sea surface elevation is affected by monsoonal effects, the difference between the two monsoons being 40 cm on the east coast of Peninsular Malaysia and 20 cm on the west coast and in Sabah and Sarawak. A tidal bore occurs at spring tide, running 40 km upstream from the entrance at Batang Lupar. Measurements of sea level using automated instruments arranged in a network of observations throughout Malaysia started in 1984, with eleven stations in Peninsular Malaysia and nine in Sabah and Sarawak. There is no discernible trend in the measured sea level variations of the 20 tidal stations around the country over the 15 years from 1984 to 1998 (Lee and Teh 2001). The Indian Ocean tsunami generated by an earthquake off Sumatra in 2004 had a major impact on the west coast of peninsular Malaysia, particularly in Langkawi (Bird et al. 2007a).

3.  Coastal Geology and Landforms The geology of Peninsular Malaysia is dominated by Palaeozoic formations and granite intrusions with a general north-south strike, while Sarawak and Sabah predominantly have Tertiary and Quaternary formations. On the west coast of Peninsular Malaysia, which is dominated by a wide coastal plain, exposures of solid rock are rare, except on offshore islands. The limited exposures on the mainland coast include Ordovician limestone at the PerlisThailand border, Cambrian red sandstone and granite along the western flank of Gunong Jerai, Silurian sediment  at the mouth of the southern bank of the Merbuk River, granites south of Perai, in Lumut-Segari, and at Minyak Beku, and granites and Devonian sediment along Negeri Sembilan-Melaka. The islands of Penang, Pangkor, Sembilan, and Besar are of granite, while Langkawi has a varied geology, dominated by limestones, quartzite, and granite. On the east coast of Peninsular Malaysia exposures of rocks are more common, although there is none along the Kelantan coast. The larger exposures include Carbon­ iferous conglomerates at Bt. Keluang; Carboniferous metasediment at Merang, Batu Rakit, and Marang; granites or Carboniferous sediment along the ­Dungun-Kuantan offset coast; Tertiary basalts at Batu Hitam; and Permian metasediment and Cretaceous-Jurassic conglomerates along southeastern Johor.

Malaysia – Introduction

The offshore islands of Perhentian, Redang, Bidong Laut, Babi Besar, and Tioman are either wholly or predominantly granitic, whereas Kapas, Seri Buat, Tinggi, and Sibu are of sedimentary rock. Rocks outcropping along the Sarawak coast, dominated by Lupar Bay and the deltas of Rejang and Miri, are rare, the few exposures being the Jurassic-Cretaceous sediment adjoining the Indonesian border, Tertiary sediment of Santubong, and the Late Tertiary sediment of Bako, Kidurong, and Tanjung Lobang. The geology of Sabah has many coastal exposures. They include Tertiary sediment of the Klias and Kudat peninsula and Sandakan, Triassic sediment and pyroclastic deposits at Lahad Datu, uplifted Quaternary corals at Semporna, and extrusive rocks at Tawau. There are great contrasts in the coastal landforms of the west and east coast of Peninsular Malaysia and those of Sarawak and Sabah. In general, the west coast of Peninsular Malaysia is sheltered from strong wave action and dominated by a wide coastal plain fringed by mangroves, whereas the more exposed east coast consists of a wide permatang (sandy beach ridges) plain, interrupted by sections of offset coasts where solid rocks outcrop to form a series of headlands. The mainly straight coast of Sarawak is dominated by the wetlands of the Rejang delta, whereas the Sabah coast is highly indented with deep bays and prominent headlands, often fringed by coral reefs. The larger rivers in Malaysia discharge large quantities of water and sediment to the sea, and only their lower reaches are estuarine. Most rivers have shoals of sand and finer sediment that move downstream, especially during episodes of flooding. In the past, rivers draining areas where there were tin mining operations have sediment yields augmented by tailings, and in some cases these have been washed down to the coast to augment beaches close to river mouths, as at Kuantan. Most tin operations have ceased in Malaysia. In Terengganu and Kelantan several lagoons have been enclosed behind sandy barriers and beach ridge plains, but infilling has been rapid, and many have given place to swampy areas. One of the larger coastal lagoons is at Tumpat, behind spits that have grown westward from Sungai Kelantan to link up with the mainland coast. Deltas and coastal plains are extensive on the coasts of  Peninsular Malaysia (Langat-Kelang confluent delta, Pahang cuspate delta, Kelantan fan-shaped delta) and Sarawak (Rejang blunt delta, Baram cuspate delta), where progradation continues around the mouths of the various rivers that deliver sand, silt, and clay to the sea. Prograding delta shores are commonly occupied by mangroves and fringed by tidal mudflats.

19.2

4.  West Coast, Peninsular Malaysia Along the west coast of the peninsula, a plain up to 50 km wide backs the coastline. A mixture of marine and alluvial accumulation, in places more than 100 m thick, has built up this coastal plain. Progradation has been accompanied by mainly mangrove advance (West Johor) and delta growth (Langat-Kelang), with occasional contributions by chenier (Kedah) and permatang development (Dindings, Lekir, Seberang Perai). The permatang of Seberang Perai lies about 2.6 m above mean sea level, and that of Beruas 2.5–5.5 m above mean sea level. An interesting feature on the west coast is the presence of well-formed relict permatang, now separated from the sea by mangroves. These permatang ridges, which represent old beaches, indicate a change in the sedimentary environment from sand to mud. A core below Permatang Sintok in Province Wellesley revealed the predominance of mangrove pollens from 1.5 to 6.0 m below the surface. Peat materials about two metres below the farthest inland ridge at Kampung Datok, Kepala Batas, were dated at 6,472 years bp (Kamaludin 2001). The base of the permatang lies about 2 m above present level, indicating that the ridge formed at a higher sea level. Towards the north, in Perak and Kedah, clearly defined terraces are found. Paramanathan and Soo (1968) recognised three levels in Perlis/Kedah behind the modern flood plain: a seasonally flooded terrace at less than 15 m, succeeded inland by 15–50 m weakly dissected terrace, and a moderately dissected terrace at more than 50 m behind it. These terraces are of Pleistocene or pre-Pleistocene origin. Granites outcrop on or near the coast at Batu Pahat, Pantai Remis, and parts of Melaka. Devonian sediment outcrop at Port Dickson and form a characteristic lateritic ironstone coast, while the plunging cliffs on the northwest of Langkawi are cut in Cam­brian quartzite. Permian crystalline limestone and Silurian-Ordovician limestone form a spectacular drowned karst coast on Langkawi. Major coastal indentations are found in the Merbuk and Dinding estuaries, but there are no major coastal protrusions, although the Selangor coast is dominated by the Kelang-Langat confluent delta. A large portion of the polder land in Perak, Selangor, and Johor is extremely low-lying, about 2.0 m below the highest tide level. These polder lands are difficult to defend, and large areas in Bagan Datoh have been abandoned.

5.  East Coast, Peninsular Malaysia The east coast of Peninsular Malaysia has a narrow coastal plain except for the Kelantan and Pahang deltas. The

1119

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19.2

Malaysia – Introduction

Kelantan fan-shaped and Pahang cuspate deltas form the two major protrusions, and central Terengganu and the southern part of the Johor coast consist of a series of alternating headlands and bays, forming a characteristic offset coastline. In the Desaru area of southeastern Johor there are many coastal features related to former high sea levels: these include high level coral heads and beach rock and oyster beds, as well as stranded beach gravels. The Desaru coast is also one of the few sites along the mainland coast where living corals still thrive. Micro-atolls are common here. Behind the beach along most of the east coast of Peninsular Malaysia is a wide belt of permatang, often consisting of distinct younger and older beach ridges, as in the area of the Kelantan-Terengganu border in Kuala Besut (Teh 1980) (>Fig. 19.2.2). Low dunes fringe the Kg. Kempadang coast, and permatang capped with aeolian sands can be observed along the S. Ular, Kijal, and Ibai coasts. Permatang ridges show ⊡⊡ Fig. 19.2.2 Permatang system of Kuala Besut.

varying heights, related to the local wave energy regime. Those in Terengganu are the highest, often cresting at about 8 m, with the youngest ridge often also the highest within the younger series, thus forming a natural bund. Beach ridges of the older series are often 3 m higher than those in the younger series. Lagoons are found in Setiu, Merang, and Tumpat, and have potential tourism value. Basalt outcrops in Beserah and Bt. Keluang is formed of conglomerates, while headlands forming the offset coast are mainly cut in metasediment.

6.  East Malaysia Three main types of coastal landforms can be recognised in Sarawak/Sabah: the broad coastal plains of Sarawak, the highly indented coast of Sabah facing the Sulu Sea, and the straight coastline with a narrow coastal plain of eastern Sabah.

Malaysia – Introduction

In Sarawak massive deltaic deposition at the mouths of the Rejang and Baram rivers has resulted in rapid progradation. The coastal plain of Rejang is more than 100 km wide, and the Baram and Trusan have been building out at an average of 10 m/year during the last 4,500 years. In these areas rapid sedimentation will partly offset the effects of future sea level rise. Permatang deposits occur in Miri, south of Kuala Sibuti, and in the Bintulu area. Spectacular sandstone cliffs are found on Bako Peninsula and Tanjung Lobang, and a cliffed coastal terrace lines the coast to the south of Miri. The mouth of the Miri River has been deflected southwards by a spit whose sand is supplied by the Baram River. Land reclamation has completely changed the shape of the Miri spit. A new channel has been dredged across the neck of the spit and the old river mouth has been filled. The west coast of Sabah facing the Sulu Sea and the Sulawesi Sea is composed of a series of deep bays with extensive swampy deltas, whereas the east coast facing the South China Sea is mainly linear, with minor indentations and a narrow coastal plain.

7.  The Malaysian States 7.1.  Perlis Almost the entire coast of Perlis is swampy and low-lying, and protected by bunds. The only area where solid rock abuts the coast is near the Thai border, where closely jointed Devonian limestone is found. Mangroves front the karsts. ⊡⊡ Fig. 19.2.3 Shelly beaches front some of the bunds protecting the swampy coastal plain of Perlis.

19.2

The coast south of Kuala Perlis is a featureless low-lying plain averaging about 1.5 m in height. The former mangroves and Melaleuca swamps have largely been replaced by settlements and agriculture. A thin belt of mangroves, interrupted by stretches of shelly beaches, persists in front of some of the bunds (>Fig. 19.2.3). The shelly beaches are recently formed, and appear to become more common in places where mangroves have been cut down or destroyed by erosion. It is supposed that the materials for beach formation comes from shells, released when the mudflat and terrace in front were eroded, lowered, and cut back. The shells are deposited as a veneer on the mud platform, and are transported inland when storm waves wash the shells across the coastal road built along the bund.

7.2.  Kedah The coastal landforms of Kedah are dominated by the featureless alluvial plain of the mainland coast and the exciting seascapes of steep coasts and drowned karsts of the Langkawi Islands (>Fig. 19.2.4). The mainland coast has rapidly prograded through the advance of mangroves, chenier development, and delta growth since the end of the Holocene marine transgression. In the Merbuk estuary, spit elongation partly enclosed the inlet and caused accelerated sedimentation and mangrove extension within the estuary and in the process tied former offshore islands to the mainland coast. Langkawi has varied coastal landscapes related to the different rock outcrops along the coast as well as a result of its island setting, which invariably shows contrasting landforms between the sheltered

1121

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Malaysia – Introduction

eastern and southern coasts compared to the more exposed western and northern coasts. The distribution of drowned topography is associated with limestone outcrops, and the prograded permatang west coast evolved under conditions of abundant sand supply and suitable wave climate. A long breakwater emplaced at the northern portion of Pantai Cenang deflected and channelled the 2004 Boxing Day tsunami waves towards Kampung Teriang, causing loss of property (Bird et al. 2007b).

7.3.  Penang Granitic Penang Island has a predominantly steep coast bordered by an apron of core stone boulders (>Fig. 19.2.5). However, the east coast is dominated by mangroves, with relict permatang behind them. The mainland coast is a low-lying plain; the exposed northern portion is composed of a permatang-chenier plain, whereas the sheltered southern portion is swampy and fringed with mangroves. A large portion of the coastal landscape of Penang has been modified by large-scale land reclamation for urban expansion.

7.4.  Perak The 230-km-long coastline of Perak, excluding Pulau Pangkor, is mainly fringed with mangroves, its inner margin defined by a coastal bund, often armoured, that protects agricultural land occupying former swamps. In Bagan

Datuk the bunds have failed, resulting in large areas of polders being reclaimed by the sea (>Fig. 19.2.6). The coastal plain is composed of a modern swampy plain built up of marine sediment, succeeded inland by a Pleistocene terrace of terrestrial sediment. Granites outcrop in Dindings, and also formed the islands of Pangkor and Pangkor Laut. Major deposits of Holocene relict permatang, linked to the nearby granite hills of Segari and Bukit Batu Tiga, occurs in Beruas and Lekir. At Matang Gelugor is a small remnant permatang, which can be traced to the Beruas system. Pleistocene terraces formed the inner margins of the coastal plain in the northern and southern parts of Perak. Along the Dindings area, the broad terrace comes right to the coast to link up with the granite headlands. A deep indentation forms the blind Dinding estuary.

7.5.  Selangor The 213-km-long Selangor coast from Sungai Bernam to Sungai Sepang can be divided into the Sekinchan coast and  the Kelang-Langat confluent delta. The delta, composed of mangrove islands protruding prominently into the Straits of Melaka, dominates the Selangor coast. The Sekinchan sector of a featureless plain is generally straight, and heavily bunded to protect oil palm plantations and paddy fields. The Selangor coastal plain was once wholly fringed by mangroves. However, intensive development for agriculture, aquaculture, and industry has resulted in large conversions of mangroves, especially in the delta. There are no rock outcrops along the coast, and any ­protrusions are

⊡⊡ Fig. 19.2.4 Cliffs cut in limestone on Langkawi Island.

Malaysia – Introduction

19.2

⊡⊡ Fig. 19.2.5 Granitic coast of Penang Island.

⊡⊡ Fig. 19.2.6 Bund failure has resulted in drowning of agricultural land.

depositional in nature. New shell beaches or even narrow beach ridge plains have developed naturally in recent years as a result of the retreat of the shell-rich marine mud terrace, and developed as recreational beaches. A new beach was created by deliberately clearing mangroves at Bagan Lallang, where a wide intertidal sand flat of ridge-runnel is exposed during low tide (>Fig. 19.2.7).

7.6.  Negeri Sembilan The coast of Negeri Sembilan (including Cape Rachado, which is part of Melaka) is about 60 km long. Despite its short length, the coastal landforms are highly varied. Rocky coasts, usually fronted by a laterite platform extending into the sea, form 24% of the coastal landforms,

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⊡⊡ Fig. 19.2.7 Ridge and runnel sand flat at Bagan Lallang.

beaches form 33%, and the rest are mangroves. The coastal geology is very uniform: Devonian folded and partly meta­ ­morphosed argillaceous rocks abut the coast along Port Dickson and part of Pasir Panjang Bay to form headlands and bluffs. Within the weathered soil profile are massive and nodular laterites, which on exposure along the coast form ironstone shore platforms with elevations from above the high water mark to below the low tide mark (>Fig. 19.2.8). Quaternary sediment of marine and alluvial origin dominate Lukut Bay, and sedimentation continues to be active as evidenced by mangrove advance. Towards the Melaka border active sedimentation continues in a mangrove environment, but a major portion of the mangrove coast is being converted. Relict beach ridges are found in Magnolia Bay and along Pasir Panjang Town.

7.7.  Melaka The 73-km-long Melaka coast has a rhythmic form, with the northern portion composed of headlands alternating with bays to form an offset coast, and the southern portion composed of a series of shallowly scalloped bays with protrusions spaced at about 4.5 km. The northern sector has a narrow coastal plain with sedimentary rocks or granites outcropping along the coast; the narrow plain is eroding, and groynes have been emplaced to slow down shoreline retreat (>Fig.  19.2.9). This con­­ trasts with the sheltered southern coast, dominated by mangroves and sandy and muddy inter-tidal flats. Narrow permatang ridges occupy a narrow belt along the entire

coast. At Kampung Pasir is an older series trending inland, and a younger series trending parallel to the present coastline. Behind Melaka Town sediment and recent alluvium fill narrow river valleys and the flat coastal plain. Granites outcrop at Tanjung Batu Supai, Tanjung Bidara, and Tanjung Ketapang. The coast near the town has been largely modified by a series of land reclamations, using materials from the sea floor and from weathered rocks. Although laterites are common in the hinterland, there is only a small coastal exposure of ironstone at Tanjung Keling.

7.8.  Johor Johor, with 492 km of coastline on the west, south, and east coast of Peninsular Malaysia, has varied coastal landforms. Mangroves wholly fringe the sheltered west coast; the south coast is characterised by drowned river valleys and broad estuaries; and the east coast is rocky, interrupted by bays arranged in an offset pattern in the south, a linear rocky coast in the centre, and a tombolo coast in the north. The broad coastal plain of mainly peat swamp forests dominates the west coast in contrast to a general absence of a coastal plain, except behind Teluk Mahkota on the east coast. The south coast is predominantly composed of terraces, with little recent alluviation. Along the west coast granite outcrops in Batu Pahat, and a relict permatang plain lies in front of the outcrop and extends southeastward before tapering away. There is no active permatang formation along the west coast. Much of the mangrove and peat swamp has been drained and

Malaysia – Introduction

19.2

⊡⊡ Fig. 19.2.8 Dissected ironstone shore platforms at Port Dickson.

⊡⊡ Fig. 19.2.9 Flat, sandy plain fronted by beaches interrupted by groynes along the northern coast of Melaka.

converted to agriculture. In places the mangrove fringe has been lost, and waves now attack the earth bunds. In response, bunds have been armoured, or a new and stronger bund built farther inland in a form of managed retreat (Teh and Bird 1999). Shoreline erosion has changed the coastal morphology, with extensive development of microcliffs and shelly beaches. The mainly mangrove-fringed south coast is drained by Sungai Pulai, Sungai Skudai, and Sungai Johor, each of which has an estuarine mouth. There is little ­sedimentation

along Sungai Johor, which still has a broad 8-km-wide mouth, and tidal influence extends 40 km inland. The pegmatite headland at Kampung Gambut, which forms the eastern end of the southern coast, has been modified by bauxite mining and small-scale land reclamation. At Kampung Gambut the coast is rugged, with small bays where there are gravel beaches up to boulder size. Behind some of the gravel beaches are high-level terraces formed of old gravel deposits with layers of shell remains. Two of the shell beds lying 2.3 m and 1.94 m above mean sea level

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were radiocarbon dated at 4120 years bp ± 65 and 2121 years bp ± 66, respectively (Bird et al. 2007a). The offset coast to Jason Bay is characterised by headlands of Permian metasediment with a truncated surface. Some headlands have rocky shore platforms with thriving coral communities and micro-atolls, while others are strewn with stranded dead coral heads. Teluk Mahkota is a semicircular bay backed by a wide permatang plain, the sands of silica sand quality. The Mersing coast is a combination of rocky headlands with shore platforms occasionally backed by a high gravel terrace and sandy beaches. The tombolo coast is a series of rocky headlands and islands, some tied to the mainland by sandy deposition. North of Kempit headland, a north elongating spit has recently developed to trap a lagoon, and mangroves are thriving in the sheltered portion of the lagoon (>Fig. 19.2.10).

7.9.  Pahang The coastline of Pahang, including Tioman Island, is about 271 km long. Long stretches of sandy beaches dominate the mainland coast, and Tioman Island is rocky. Three major coastal sectors can be recognised in Pahang: the southern spit coast, the central delta coast, and the northern offset coast. The spit coast from Endau to Tanjung Batu is dominated by a series of southward elongating spits at the mouths of the Endau, Pontian, Rompin, Mercung, and Bebar rivers (>Fig. 19.2.11). The spits, which run parallel to the coast, have a complex history of changing outlets and forms (DID 2006).

The delta coast extends from Tanjung Batu to Sungai Kuantan. The southern part of the cuspate delta, occupied by extensive peat swamps, differs greatly from the sandy northern portion of beach ridges. The offset coast of headlands with intervening bays has evolved through the in­­ filling of the embayments created at the height of the Holocene marine transgression and the linking of offshore islands (DID 2002). Tioman Island has a mainly rocky coast, often with steep cliffs. Some of the rocky granite isles off Pahang show well-developed fluting (>Fig. 19.2.12).

7.10.  Terengganu The coast of Terengganu is composed of an offset coast of prominent headlands with intervening bays in the south, a central linear coast of permatang plain, and to the north are the Setiu-Merang lagoon and the cuspate coast of welldeveloped and well-preserved beach ridges. The beach ridge pattern behind the conglomerate Bukit Keluang headland has a complex pattern, recording the evolution of the coastline, which tied the former Keluang Island to the mainland.

7.11.  Kelantan The Kelantan coast, which is wholly sandy, forms the outer edge of the Kelantan coastal plain, built up by deltaic, marine, and swamp deposits. The coast, unprotected by headlands or offshore islands, is completely exposed to waves from the South China Sea, has a concave sector at

⊡⊡ Fig. 19.2.10 Beach ridges, Kempit headland, and barrierlagoon system.

Malaysia – Introduction

19.2

⊡⊡ Fig. 19.2.11 Distal end of Mercung spit.

⊡⊡ Fig. 19.2.12 Densely fluted granite on Labas Island.

Sungai Semerak, a protrusion at Pengkalan Datu, and an irregular coastline of deltaic islands, spits, and a lagoon at Tumpat. The low-lying coastal plain is very wide, composed of a 10-km outer belt of barrier and deltaic deposits backed by a 30-km-wide alluvial plain whose surface is often interrupted by abandoned levees and meander scrolls. The coastal alluvium is deep, attaining more than 100 m along the outer edge of the plain.

7.12.  Sarawak The Rejang delta, formed of wetlands, dominates the mainly straight coast of Sarawak. However, delta formation is common where progradation continues around the mouths of the various rivers that deliver sand, silt, and clay

to the sea. Prograding delta shores are commonly occupied by mangroves, and fringed by tidal mudflats. Massive deltaic deposition at the mouths of Rejang and Baram rivers has resulted in rapid progradation. Beach ridge deposits occur in Miri, south of Kuala Sibuti, and in the Bintulu area. A new river mouth has been cut along Miri spit, and the old entrance filled in and reshaped by land reclamation. Spectacular sandstone cliffs are found on Bako Peninsula and Tanjung Lobang, and south of Miri the beach is backed by a coastal terrace of partially consolidated sediment.

7.13.  Sabah The southeast coast facing the Sulawesi Sea is made up of Tawau Bay, Semporna Peninsula, and Lahad Datu Bay.

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Tawau Bay is mainly a mangrove coast, with extensive mangrove islands around the head of the bay. Semporna Peninsula is composed of mainly dacite and andesite flows along the south-facing coast, uplifted corals along the eastern tip, and extrusive rocks and a Jurassic–Cretaceous chert-spillate formation along the northeast-facing coast. Lahad Datu Bay is a low-energy coast further protected by fringing corals, and rivers discharging into the bay have formed deltas at the mouths of the Tingkayu and Matamba rivers. Lahad Datu Bay is believed to have subsided, and has a relatively narrow coastal plain. The northwest coast facing the Sulu Sea is a series of deep bays with extensive swampy deltas, including the Dent Peninsula, Eastern Deltas, Sandakan Bay and Peninsula, Labuk and Sugut Deltas, Bengkoka Peninsula, and Marudu Bay. Deltas have formed at the mouths of all the major rivers flowing into the Sulu and Celebes seas. Deltas of the Kinabtangan, Segama, and various smaller streams have coalesced, forming about 1,800 sq km of low-lying swamps that Collenette (1963) described as the Eastern Deltas. The Labuk and Sugut deltas are composed of mangrove islands separated by a network of tidal channels. Mangroves and nipah dominate the Labuk Delta, with many wide tidal rivers discharging muddy water into the bay. The Labuk Delta lies within the Labuk Bay in contrast to the Sugut Delta, which protrudes boldly into the Sulu Sea. There are sandy beaches and relatively large rivers 5–10 km apart. The Bengkoka coast has rocky headlands, with minor bays and mangroves in sheltered sectors. There are long stretches of rock exposure of various ages on both sides of the peninsula. Marudu Bay is a sheltered mangrove-fringed coast, with small patches of fringing corals persisting despite the murky water. The west coast of Sabah includes the Klias Peninsula, Papar Delta, Tuaran-Kota Kinabalu coast, Kota Belud coast, and Kudat Peninsula. The Klias Peninsula is a former ria coast in which the inlets have been filled by active sedimentation during the Holocene. Here, beaches and spits line Kimanis Bay, the sands supplied by the Papar River to the north. The coast fronting the South China Sea consists of retreating cliffs with a narrow beach, and the southwestern corner of the peninsula is formed of deltaic deposits colonised by mangroves. The cuspate Papar Delta consists of permatang deposits in the southern arm and an elongated spit along the northern arm. The Tuaran-Kota Kinabalu coast is highly diverse, with indentations in Usukan, Am­­ bong, Sulaman, Mengkabong, and Sepangar bays, deltas in Kinabatangan and Tuaran, and barrier spits in Karambunai

and south of Kota Kinabalu Town. The Kota Kinabalu coast has been extensively modified by land reclamation. Extensive low-lying permatang deposits are found on the Karambunai Peninsula, and coral reefs are common near Pulau Gaya. The Kota Belud bay is shallowly indented, with a slight concave northern end, and sandy beaches backed by a permatang plain up to 4 km wide. The Kudat coast is still actively evolving, as small bays are filled, offshore islands tied to the mainland, and cliffs cut back. Fringing coral reefs protect the whole of the east coast of Kudat Peninsula, and extend midway into Marudu Bay. The Sikuati coast consists of a 4-km-wide permatang plain, in contrast to the narrow permatang plain resting on a Pleistocene marine platform between Kudat and Tanjung Agong Agong.

References Bird MI, Teh TS (2006) Mangroves and urbanisation in Southeast Asia. In: Wong TC, Shaw BJ, Goh KC, (eds) Challenging sustainability: urban development and change in Southeast Asia. Marshall Cavendish, Singapore, pp 3–52 Bird MI, Fifield LK, Teh TS, Chang CH, Shirley N, Lambeck K (2007a) An inflection in the rate of early mid-Holocene sea level rise: a new sea level curve for Singapore. Estuar, Coast Shelf Sci 71:529–536 Bird MI, Cowie S, Hawkes A, Horton B, McGragor C, Ong JE, Tan A, Teh TS, Zulfligar Y (2007b) Indian Ocean tsunamis: environmental and socio-economic impacts in Langkawi. Malaysia. Geogr J 173:103–117 Collenette P (1963) A physiographic classification of North Borneo. J Trop Geogr 17:28–33 DID (Drainage and Irrigation Department) (2002) Integrated shoreline management plan of Northern Pahang, Government of Malaysia, Kuala Lumpur DID (Drainage and Irrigation Department) (2006) Integrated shoreline management plan of Southern Pahang and Tioman, Government of Malaysia, Kuala Lumpur EPU (Economic Planning Unit) (1985) National Coastal Erosion Study Lee SC, Teh TS (2001) Coastal resources. In: Chong AL, Mathews P (eds) Malaysia. National response strategies to climate change. Moste, Malaysia Kamaludin H (2001) Geological evidence: Core samples. In: Ong JE, Gong WK (eds) Encyclopedia of Malaysia, vol 7. Seas, Archipeligo Press, Singapore Paramanathan S, Soo SW (1968) Reconnaissance soil survey of Kedah. Malayan Soil Survey Report 3. Ministry of Agriculture and Co-operative Teh TS (1980) Morphostratigraphy of a double sand barrier system in Peninsular Malaysia. Malaysian J Trop Geogr 2:45–56 Teh TS, Bird ECF (1999) Managed retreat: with special reference to the mangrove coast of Sungai Lurus in Peninsular Malaysia. Malaysian J Trop Geogr 30:39–50 Teh TS, Lim CH (1993) Impacts of sea level rise on the mangroves of Peninsular Malaysia. Malaysian J Trop Geogr 24:57–72

19.3  Singapore

Wong Poh Poh

1.  Introduction Singapore is a city-state located at the southern end of the Malay peninsula. Large-scale landfill projects starting from the 1960s have substantially changed its coastline (Wong 1985), which has advanced seaward by 10 km at its western end (>Fig.  19.3.1). This modified or developed coastline has significant harbours, docks, and various coastal protection structures, while barrages have been built to enclose estuaries along the west coast, north coast, and at the Marina channel. Along the east coast and at Pasir Ris the modification has included the forma­­ tion  of  beaches between series of breakwaters acting

as ­headlands, thus producing a sandy coastline. Rock bunds were constructed on the reef platforms around many of the southern islands, with gaps for beaches to form. The tides in Singapore are semidiurnal and of a mixed type, with large diurnal inequalities in the low tide levels. The tide range is 2.6 m. The climate is hot and humid, and conducive to rapid weathering and other subaerial processes. Despite the protected wave environment, the southern islands of Singapore had a wide variety of coastal features (Swan 1971). Where the coast has not been modified, the influence of geology is still present, for example the granite corestone boulders at Changi and on Pulau

⊡⊡ Fig. 19.3.1 Singapore.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 19.3.2 Granite tors with pseudokarren on the coast of Pulau Ubin.

⊡⊡ Fig. 19.3.3 Chek Jawa covered by a low tide.

(Island) Ubin, and Triassic sedimentary rocks forming cliffs, steep slopes, and limited gravel beaches on the west coast and on some southern islands. Mangroves are limited to one major area in the northwest, Pulau Tekong and Pulau Ubin. Most of the fringing coral reefs in the western half of the island and around the southern islands have been destroyed by landfill projects.

2.  The Singapore Coastline Pulau Ubin is dominated by Lower mid-Triassic granite, and has four distinctive shore types: coastal tors as headlands (>Fig. 19.3.2); low cliffs of weathered granite; short stretches of sandy beach, with many overlain by granite gravels from quarries; and mangroves. The weathering

Singapore

19.3

⊡⊡ Fig. 19.3.4 Changi spit.

⊡⊡ Fig. 19.3.5 Beach compartments between headland breakwaters on the reclaimed east coast of Singapore.

of granite boulders has produced solution grooves or pseudokarren, especially at the eastern end (Tschang 1961). Immediately north of Tanjung (Cape) Chek Jawa is a lowtide lagoon (>Fig.  19.3.3) formed behind a cuspate sand bar and harbouring several coastal ecosystems. Pulau Tekong retains its low cliffs and steep slopes of sedimentary rocks, mangroves, and small beaches on its northern coast. The only sand spit in Singapore is at the western end of Changi Beach (>Fig.  19.3.4). With a length of about 150 m, the spit provides a sheltered area in the river for small boats and a ferry terminal. Changi Beach, until recently the longest remaining stretch of natural beach,

has been enlarged through beach nourishment for coastal recreation. The reclaimed east coast has a sandy shore that provides several interesting aspects of the concept of deploying breakwaters as headlands for intervening beaches to form (>Fig.  19.3.5), imitating zetaform bays in nature (Wong 1981). The landfill material for this coast was derived from a predominantly sandy Quaternary deposit that forms most of the eastern part of Singapore. A seawall was built initially to protect the reclaimed coast, but this did not encourage beach formation. Subsequently, two types of breakwaters acting as headlands were constructed.

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⊡⊡ Fig. 19.3.6 A tombolo formed behind a headland breakwater, east Singapore.

⊡⊡ Fig. 19.3.7 Sand washed up over pneumatophores of mangroves at Pasir Ris.

Small gabion breakwaters were built on the foreshore of the fill, and were subsequently enlarged and increased in height to have riprap surfaces. Large riprap breakwaters were constructed dry within the fill. As the fill is removed by marine action, isthmuses are formed. With the subsequent removal of the isthmuses, salients and tombolos are formed behind the breakwaters (>Fig.  19.3.6). Basically, the beach along the reclaimed east coast is the sand derived from the fill material and overlies the sloping but irregular

surface of the fill in the intertidal zone. At Pasir Ris a series of riprap breakwaters were also deployed to protect the reclaimed land, but beach formation has been limited due to the low wave energy in the Straits of Johore. Mangrove saplings washed out from nearby rivers have colonised two bays and helped to form beaches, a situation not found on the east coast (>Fig. 19.3.7). Extensive reclamation projects occurred off the southwest coast of Singapore, including the forming of artificial

Singapore

19.3

⊡⊡ Fig. 19.3.8 Artificial beach at Sentosa.

⊡⊡ Fig. 19.3.9 Cliff at Tanjung Rimau, Sentosa.

islands and the enlargement of natural islands. Jurong Island is the amalgamation of nine main islands. Much of the western coastline of Sentosa has been modified to form artificial beaches (>Fig. 19.3.8). Its southern coastline has been reclaimed for a marina and waterfront properties, although some cliffs remain at its northern headland (>Fig. 19.3.9).

The west coasts of St. John’s Island and Lazarus Island remain as natural cliffs of conglomerate and sandstone, while their east coasts have been modified by stone bunds enclosing beaches. Lazarus Island is linked to the reclaimed island of Pulau Seringat to the north. At the southern end of St John’s Island is a fine anticlinal fold, with its axis aligned northwest–southeast, and dips 40° SW and 55° NE

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⊡⊡ Fig. 19.3.10 Replanted mangroves at Pulau Semakau.

(Public Works Department 1976). Pulau Tekukor has the same geology, and its almost vertical beds have resulted in cliffs with virtually no shore platform or beach. Pulau Subar Darat, Pulau Subar Laut, and Kusu Island are typical examples of rock bunds constructed to surround the island core and the area enlarged by sand dredged from the sea to form beaches in the gaps between the structures. A 7-km rock bund enclosed the shallow area between Pulau Semakau and Pulau Sakeng to form an offshore sanitary landfill site. Replanted mangroves have flourished on Pulau Semakau outside the perimeter bund (>Fig. 19.3.10). Pulau Biola, consisting mainly of Triassic sandstones, remains as a rocky island smoothed by subaerial and coastal processes. In contrast, the cliff outline on Pulau Salu is a more irregularly stepped profile because of silicified tuffs resistant to subaerial weathering (Swan 1971).

The west coast of Singapore retains its cliffs but the mangroves have disappeared as a result of the construction of barrages. Beaches are absent on this low-energy coast. At the northwestern end is Singapore’s major mangrove park with its mature mangrove trees and intertidal mudflats.

References Public Works Department (1976) Geology of the Republic of Singapore. Singapore Swan SB St. C (1971) Coastal geomorphology in a humid tropical low energy environment: the islands of Singapore. J Trop Geogr 33:43–61 Tschang HL (1961) The pseudokarren and exfoliation forms of granite on Pulau Ubin. Singapore. Z Geomorphol 5:302–312 Wong PP (1981) Beach evolution between headland breakwaters. Shore Beach 49:3–12 Wong PP (1985) Artificial coastlines: the example of Singapore. Z Geomorphol 57:175–192

19.3.1 Thailand: Gulf of Thailand Coast

Sanit Aksornkoae · Eric Bird

1.  Introduction The east coast of peninsular Thailand is dominated by a wide depositional lowland, interrupted by ridges that run out from the interior mountains to protrude as headlands (Pitman 1982). Along the Gulf coast, as far north as Petchaburi, Tertiary sediment are confined to small basins, and Tertiary orogenic movements resulted in the subsidence of the Chao Phraya depression, north of the Bight of Bangkok. The Gulf of Thailand began to form as a result of late Cretaceous to early Tertiary subsidence, which emphasised the north-south tectonic trends, and formed a series of pronounced depositional basins occupied by Quaternary marine sediment. Tjia et al. (1977) reviewed Quaternary sea level changes in this area, and concluded that there is good evidence that the sea rose temporarily above its present level in Holocene times. There are several emerged coastlines, marine terraces, and river terraces indicative of the recent northwest tilting of the peninsula and the smooth emergent coast of the Gulf. Shallow offshore gradients ( Fig. 19.3.1.1), while Hinggam Beach consists of coarse, yellow quartz sand contained between two forested capes of weathering granite. The coast turns westward past headlands of Permian limestone, and Ko Samui and Ko Pha Nga are high forested islands offshore. Ban Don Bay has extensive tidal mudflats backed by a broad mangrove area. At Surat Thani the birdsfoot delta of the Mae Nam Ta Pi branches through mangroves, with

⊡⊡ Fig. 19.3.1.1 The wide prograding Savaniwat Beach, on the Gulf of Thailand coast north of Nakhon. (Courtesy Geostudies.)

Thailand: Gulf of Thailand Coast

frequent channel switching. The river has a high sediment load, and wave energy is low. On the northern side of Ban Don Bay, Laem Si is a shelly sand spit that is growing southward in a sector where southeasterly wave action is excluded by the mountainous promontory to the south. It has recurved ridges curving back into mangroves. To the north sandy beaches extend between rocky promontories, and there is a steep coast sector at Ko Wiang. The coast, for a distance 310 km north of Laem Si, has long stretches of sand beach, backed by a rapidly narrowing coastal plain, and broken only by the south–southwest/ north–northeast trending ridges of Cretaceous and Triassic sandstone and limestone, which form headlands and offshore islands between Lang Suan and Prachuap Khiri Khan. The sheltered bays inshore of these islands are commonly filled by beach ridges and lagoons, with bay head mudflats and mangrove swamps. Pak Nam Chumpon is a seaside resort, with jagged limestone islands offshore, and beaches that include sand eroded from cliffs of weathered sandstone. Near Prachuap Khiri Khan there are bays with curving beaches backed by beach ridge plains and tombolos attaching islands as headlands, as at Khao Lom Muak, as well as on several high islands (Ko Lak, Ko Raet). Ao Noi is a small bay between attached headlands. The tombolos typically have parabolic sand beaches, swamps in the lee of the island, and major streams emerging in their lee, as at Laem Mae Ramphung. The steep Khoo Sam Roi Yot National Park coast is to the north, and a bouldery granite headland is followed by

⊡⊡ Fig. 19.3.1.2 Eroded beach at Laemlaut, near Phetchaburi. A sand ridge has moved in across a mangrove swamp, fragments of which remain on the shore. (Courtesy Geostudies.)

19. 3.1

Ban Pak Nam Pram beach, with a large Buddha statue in the sea. From Hua Hin seaside resorts (Majestic Beach, Royal Garden Village, Tocanee Beach, Praramhok, Golden Sands, The Regent) line beaches northward, and there are longshore spits nourished by drifting in that direction. The Had Petch beach resort has been narrowed by erosion, countered by building a sea wall and stepped groynes. The Phetchaburi River has built a large delta, with former mouths marked by three cuspate forelands, backed by tidal swamps and old beach ridges. On one of these forelands the spit at Laem Pak Bia grew 2 km northward between 1961 and 1967. At Laemlaut a sand ridge has been driven in across a mangrove swamp by a storm surge (> Fig. 19.3.1.2), deflecting a mangrove creek northward. The beaches come to an end, and mangroves fringe the coast that turns eastward past the Mae Nam Mae Khlong Delta along the head of the Gulf of Thailand.

3.  Head of the Gulf of Thailand The coast at the head of the Gulf of Thailand is dominated by a 10-km-wide zone of tidal mudflats backed by mangrove swamps, which have been extensively modified to prawn ponds and salt works. Behind these, brackish-water swamps have been largely converted to ricefields. The Mae Kong and Tha Chin rivers meander through the swamps to the sea. The coastal plain has been formed by the combined deltaic sedimentation of major rivers, of which the

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Thailand: Gulf of Thailand Coast

Mae Nam Chao Phraya is the largest (Coleman and Wright 1976). The Chao Phraya River and its tributaries drain an area of 11,329 sq km, and have formed a flood plain that extends 470 km south from the 100 m contour to the 120-km-wide delta front (> Fig. 19.3.1.3). From Ayutthaya southward the terrain is very flat and averages 2 m above sea level, except for a major area east and north of Bangkok, 3.5–5 m high, possibly an old barrier island. Peat from the backswamp dated as 3,100 years bp suggests that the delta front has advanced at least 4 m annually. Offshore, the slope is exceedingly low ( Fig.  19.3.1.4). The eastern side of the

Gulf is flanked by smoothly curved sandy bay-head beaches and forested bluffs between rocky headlands of sandstones and granite. At low tide splays of sand and gravel are exposed, often with large ripples (> Fig. 19.3.1.5). The southeast coast between Satthip and Hiaeng has a series of three wide, sandy bays formed between mountainous ridges of igneous rocks and limestone, which here run north–south west of Ko Samet, then northeast– southwest to the east. These ridges continue offshore as steep rocky islands. The sandy beaches, of which Rayong is typical, are backed by alternating parallel ridges and depressions, with shallow lagoons in the depressions. Total relief seldom exceeds 5 m, although dunes up to 12  m high occur locally. At Rayong the beach ridges extend up to 7 km inland, and there are eroding dunes at Map Ta Phut (> Fig. 19.3.1.6). East of Klaeng broader alluvial plains have formed between the northwest–southeast trending ridges by the sediment influx from the Krasae, Phang Rat, Chanthaburi, and Welu rivers. The more sheltered mouths of the Krasae and Phang Rat rivers are mangrove-fringed, with broad tidal marshes inland between inward-curving offshore islands. The coast south of Chanthaburi is more exposed to the southwest monsoon, so that a complex pattern of  beach ridges has evolved. Behind the outer barrier and  the former coastline extensive (180 sq km) mangrove swamps and tidal marshes have developed, broken

⊡⊡ Fig. 19.3.1.3 Prawn ponds on the coastal plain, northern Gulf of Thailand.

Thailand: Gulf of Thailand Coast

19. 3.1

⊡⊡ Fig. 19.3.1.4 Beach rock ramp south of Pattaya. (Courtesy Geostudies.)

⊡⊡ Fig. 19.3.1.5 Steep coast near Pattaya showing sandy beach fronted by sand and gravel foreshore exposed at low tide. (Courtesy T. Burgers.)

­ ccasionally by isolated inselbergs, some of which form o offshore islands. From the Welu River to the mudflat-fronted mangrove swamps of Trat Bay the coast becomes rocky as shales and sandstones of the Kanchanaburi formation outcrop. Ko

Chang forms a rugged island of Tertiary igneous rocks sheltering Trat Bay. From Trat to Khlong Yai the coastal plain rapidly narrows to less than 500 m before rising toward the Banthat Range. Here longshore drifting is northward, with offshore bars and spits.

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⊡⊡ Fig. 19.3.1.6 Eroded dune, Map Ta Phut. (Courtesy Geostudies.)

References Aksornkoae S (1976) Structure, regeneration and productivity of mangroves in Thailand. Michigan State University, East Lansing, MI Bird ECF (1989) The effects of a rising sea level on the coasts of Thailand. ASEAN J Sci Technol Dev 6(1):1–13 Coleman JM, Wright LD (1976) Modern river deltas: variability of processes and sand bodies. In: Broussard MLS (ed) Deltas: models for exploration. Houston Geological Society, pp 99–149

National Environment Board (1985) Songkhla lake basin planning study. Bangkok Pitman JI (1982) Geomorphology of the coastline of Thailand. Occasional paper. Department of Geography, King’s College, London Tjia HD, Fujii S, Kigoshi K (1977) Changes of sea level in the southern South China sea area during quaternary times. In: Proceedings of the quaternary geology of the Malaya-Indonesia coastal and offshore areas seminar, 13th CCOP Session, pp 11–49, Publication 5

19.4  Brunei (Negara Brunei Darussalam)   Gabriel Yong   

1.  Introduction Brunei occupies a 130 km stretch of the northwestern coast of Borneo between longitudes 114°4' E and 115°9' E, which includes parts of the Baram Delta in its western section and the Brunei Bay in its eastern section. The continental shelf is 50–80 km wide and contains submerged sand shoals, rocky outcrops and patch reefs (> Fig. 19.4.1). The main sources of information on Brunei’s coast include the works of Wilford (1961), Tate (1970, 1971), James (1984), Chua et al. (1987), Goh (1981) and Sandal (1996). Brunei has a consistently warm climate (mean monthly temperature range 27–28°C) and is wet (annual rainfall range 2,500–4,500 mm) throughout the year because of its geographic location (4–5° N of the Equator). Fluctuation of the Inter-Tropical Convergence Zone (ITCZ) gives rise to two seasons, the northeast and southwest monsoons. Although there is high spatial and temporal variability in rainfall distribution, it is generally wetter during the northeast monsoon from November to January. Climatic conditions are also affected by cyclones in the South China Sea, which produces squally weather, as well as the El Nino Southern Oscillation (ENSO), which bring about dry spells (El Nino) and floods (La Nina). The coastal environment is controlled largely by the interaction of the country’s geology, climate and wave regimes. Increasingly, however, As over 80% of the country’s population lives along the coast, human activities have modified, and will continue to alter, Brunei’s coastal landscape through mining and  development.

2.  Coastal Environment From November to January, the northeast monsoon generates storm waves that arrive at the coast from the north and northeast with wave heights and periods typically of 1–3 m and 4–6 s respectively. Storms in the South China Sea generate destructive swell waves that scour the beach face and backshore and produce offshore sand bars. Along the Brunei-Muara coast rip currents and undertows are common, while in the Tutong-Belait coasts to the west

large waves generate strong longshore currents and a westward longshore drift. The sea is mostly calm during April and May. The wave regime during the southwest monsoon is generally constructive. Low-energy swell waves (height Fig. 19.4.3) caused by intense fluvial activity, which cuts readily into the soft sedimentary rocks, weakening cliff faces through gullying and groundwater flows. The talus accumulated at the foot of the cliffs is then removed by waves during high tides or storms. The process is accelerated by development practices that leave large areas of bare ground unprotected. Sand mining at Tungku Beach in the early 1990’s exacerbated coastal erosion. Today, large sections of the coastline in

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⊡⊡ Fig. 19.4.3 Eroding coast, Berakas.

the Brunei-Muara District, e.g. at Jerudong and Berakas, are protected by detached headland breakwaters and beach nourishment. At Penanjong, where cliff recession is most rapid, the cliff face is now protected by retaining walls. The reconstructed coastline at Tg Punyit and Tungku is protected by rock revetments. Along the Brunei-Muara coast longshore drift has formed sand barriers that dam small rivers, and sand spits that  deflect river mouths to the east, such as the Pelumpong Spit at the eastern end of the Brunei-Muara coast. Just west of Tag. Batu, a mangrove swamp has developed around Sg. Pemburongan as a result of restricted drainage cause by the barrier beach. The sheltered coast of Brunei Bay is lined with mudflats and mangroves fronting high sandstone ridges. Marine transgression during the Holocene, which drowned river valleys and the narrow continental shelf, formed the Sungai Brunei estuary and several rocky islands. Although four rivers (Imang, Damuan, Kedayan and Kianggeh) drain into Sungai Brunei, its freshwater input is relative small and the estuary is well mixed and saline for most if its 16 km length leading into the Brunei Bay. Mangrove forest lines the shores of the estuary and islands, particularly the northeastern shores in the lee of ebb currents, where conditions are conducive to mudflat development.

The Temburong coast has mangrove swamps along the southern shore of Brunei Bay. This is sometimes referred to as the Inner Bay, and is separated from the rest of Brunei Bay by the Trusan delta. The Limbang and Trusan rivers drain into the bay, bringing large quantities of sediment from catchments in Sarawak.

References Chua T-E, Chou LM, Sadorra MSM (eds) (1987) The coastal environmental profile of Brunei Darussalam: resource assessment and management issues. ICLARM Tech. Report no. 18 Goh KC (1981) The state of the physical environment of Brunei Darussalam. Malaysian J Trop Geogr 22(1):19–28 James DMD (ed) (1984) The geology and hydrocarbon resources of Negara Brunei Darussalam, Museum Brunei: Bandar Seri Begawan Sandal ST (ed) (1996) The geology and hydrocarbon resources of Negara Brunei Darussalam, Brunei Shell Petroleum, Seria Tate RB (1970) Longshore drift and its effect on the new Muara Port. Brunei Museum J 2(1):238–252 Tate RB (1971) Radio-Carbon ages from Quaternary terraces – ­prehistory in Brunei. Brunei Museum J 2(3):108–123 Wilford GE (1961) The geology and mineral resources of Brunei and adjacent parts of Sarawak with description of Seria and Miri oilfields. British Borneo Geological Survey, Brunei Government

19.5 Cambodia

Eric Bird

1.  Introduction

Ream peninsula to the southeast. The Palaeozoic rocks of the Elephant Ranges lie to the east, and beyond these the coast is low-lying towards the Vietnam border and the Mekong Delta (>Fig. 19.5.1). The coastal climate is tropical, with a wet summer monsoon (May–October) and a dry winter season. Annual rainfall at Kampoh is 1,976 mm. River flooding is extensive in the wet season.

Cambodia (also known as Kampuchea) has an embayed and island-fringed coastline about 400 km long. The western part is dominated by Mesozoic sandstone, which culminates in the Cardamom Mountains, rising more than 1,500 m above sea level. The broad Bay of Kompong Som is a low-lying synclinal region, bordered by the uplifted

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Wave action in coastal waters is generally slight, but it becomes stronger during the southwest monsoon, especially when it is reinforced by sea breezes. Typhoons from the China Sea occasionally move in through Cambodian waters, producing temporary storm surges with strong wave action. Tidal ranges are generally less than 1 m (0.4 m at Rëam), the tides being diurnal. Current action is weak, except in narrow straits (e.g. between islands), where tidal scour may be intensified. There is evidence that the sea has stood at higher and lower levels relative to the land along the Cambodian coast. Carbonnel (1972) found evidence of Quaternary marine terraces 25, 10–15, 4, and 1.5–2 m above present sea level along the higher parts of the Cambodian coastline. The lowest of these terraces is regarded as Holocene in age. Extensive submergence during the late Quaternary marine transgression formed bays and inlets, especially at valley mouths. Coral reefs are generally offshore, around islands and fringing headlands, and there are mangroves alongside estuaries and on parts of the coast sheltered from strong wave action.

2.  The Cambodian Coast Just south of the Thailand border a hilly peninsula borders the wide estuary of the Flong Klun River. The coast hereabouts is penetrated by many branching, mangrove-fringed inlets. Tidal scour is strong north of Koh Kong Island, which shelters a wide, shallow indented bay from the open sea of the Gulf of Thailand. To the south the coast to Point

Samit is steep and hilly, with rocky promontories separating sandy beaches, many of which are backed by dunes bearing scrub and woodland. Banks of sand and coral are extensive offshore, Condor Reef being the largest of the outlying coral formations. Further south, Koh Rong is the largest of several steepsided high islands off the Bay of Kompong Som. The outer coasts of these islands have cliffed promontories, fringing reefs, and sandy beaches; the inner embayments have mangroves. The northwestern part of the Bay of Kompong Som is shallow and sheltered, with very extensive mangrove areas passing into the estuaries of the Piphot and Kompong Som rivers. Beach ridges and spits have formed near river mouths. Wharves have been built at Pointe Loune to establish the port of Kompong Som (Sihanoukville). The Ream peninsula generally has low-lying coasts, bordered by reef-fringed islands and shoals, with wide, funnel-shaped estuaries flanked by mangroves. The nearshore zone is shallow and rocky. To the east the hinterland steepens to the Elephant Range, spurs from which form low promontories, and the Kos Sla River on their eastern margin has built a small delta south of the town of Kampoh. From here to Vietnam the coast is low and intricate, with many islands, shoals, reefs, and rocks in the nearshore shallows, the water sheltered by the large, high Vietnamese island of Quan Phu Quoc.

Reference Carbonnel JP (1972) Le Quaternaire Cambodgien. Memoirs ORSTOM No. 60, Paris

19.6 Vietnam

D. Eisma

1.  Introduction The Vietnamese coastline extends for about 3,400 km, and includes two large river deltas, the Mekong Delta in the south and the Song Hong (Red River) Delta in the north, formed of Holocene alluvial deposits. Much of the intervening coast is backed by a narrow plain and mountain ranges, with ridges and valleys following structures along two main tectonic axes, a dominant SE–NW axis and a subordinate SW–NE axis. The main folding and faulting lines have also influenced the general outline of the coast. A basement of pre-Cambrian and early Palaeozoic granites, gneiss, rhyolite, porphyrites, and schists dominates the southern highlands, partly overlain by Mesozoic formations. North from Da Nang there are a series of SE–NW folds and faults in Palaeozoic and Mesozoic rocks. The major mountain range of Vietnam, the Annamite Mountains, lies close to the coast, behind a coastal plain 30–50 km wide. This range forms the watershed between short rivers that flow eastward into the South China Sea and those that flow westward into the broad basin of the Mekong. Much of the coast consists of rocky promontories, beaches, and small estuaries. Coral grows along most of the mainland coast as well as around the islands – Beach Long Vi, Paracel Islands, Con Son, Tho Chau Islands (Poulo Panjang) – where there is no outflow of river water nearby. Along the coast of Vietnam and on outlying islands, such as Bach Long Vi (Nightingale Island) in the Gulf of Tonkin, the Paracel Islands, Con Son (Poulo Condor), and the numerous small islands off the west coast in the Gulf of Thailand, there are indications of emerged coastlines 1–2 m and 4–5 m above present sea level (Van Lap Nguyen et al. 2000). In some areas, as at Mui Bai Bung (Cape Ca Mau), Mui Vung Tau (Cape St. Jacques), and along Phu Quoc Island, there are indications of higher emerged coastlines at 10–15 m, and near Phan Rang at 50 m and 75–80 m. former sea levels of 1–2 and 4–5 m are indicated by deposits of shells, shell fragments, and coral, fossil oysters still attached to rocks, emerged reefs, beach rock, notches and sea caves. The deposits associated with these emerged coastlines are not consolidated or weathered into laterite.

Waves and currents along the Vietnamese coast are from the northeast between November and March during the northeast monsoon. Around Mui Bai Bung (Cape Ca Mau) they turn to the northwest in the Gulf of Thailand during this period, whereas in the Gulf of Tonkin there is a northward counter-current along the coast. In July and August, during the southwest monsoon, waves and currents are predominantly from the southwest. Along the eastern shore of the Gulf of Thailand the current trends southeasterly, then turns around Mui Bai Bung toward the northeast. In the Gulf of Tonkin, and farther south along the coast past Mui Varella, the coastal current is southward during that time. Current velocities in January are about 15–85 cm/s; in July, they are 15–40 cm/s. Monsoon winds dominate, but along the eastern shore of the more sheltered Gulf of Thailand there is often an alternation of sea breezes by day and land breezes by night, whereas along the eastern Vietnam coast the wind tends to be deflected parallel to the coast. Sea waves and swell follow the same direction as the wind. The strongest waves occur from the northeast in winter, predominantly affecting the northeast-east exposed coast from Vinh to Mui Dinh (Cape Padaran). The Gulf of Tonkin is more protected, the swell entering during the winter from the south, whereas the Gulf of Thailand is fairly sheltered from the southwest as well as from the northeast winds. Typhoons, coming from east-southeast, generate high waves and storm surge flooding, and occur mainly from May to December and are rare from January to May. They occur chiefly north of 10° N and rarely enter the Gulf of Thailand. The tides are diurnal, with a strong semidiurnal component during the equinoxes. The mean spring tide range at Mui Vu Tau (Cape St. Jacques) is 2.6 m, at Hon Nieu 1.8 m, and at Hon Dau (Red River) 2.0 m. Tidal currents are generally weak. Along the central Vietnam coast they are of little influence, increasing to 1–1.5 knots along the north coast in the Gulf of Tonkin. At the Mekong Delta the water level and the outgoing current are chiefly ­determined by the river outflow, which rates up to two knots from January to June, when water level in the river is low, and up to 4.5 knots from July to December, when river water level is high.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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2.  The South Coast The south coast consists almost entirely of Quaternary deposits, in part very recent, as at the Mekong River mouths, and some rocky promontories (granite) at Mui Vung Tau (Cape St. Jacques). Directly northeast of the Mekong River mouth, which is divided into several large branches, are the estuaries of several smaller rivers (among them Song Guam Ca, Song Saigon, and Song Be). Here, a small bay remains, which the rivers have not yet filled. In front of the Mekong River mouth the coast bulges seaward and is formed of mudflats and mangrove swamps. The Mekong river delta consists of a large flat floodplain with levees and a backswamp behind a coastal plain with relict beach ridges and sand dunes (>Fig.  19.6.1). Along the coast a usually small belt of mangroves and a few sandy spits grade seaward into tidal flats 1–3 km wide. In the delta 10–20 m of Holocene (mainly deltaic) deposits overlie late Pleistocene deposits, older bedrock, and a filled-in incised river valley (more than 70 m deep), which dates from the last glacial period. In the earliest Holocene deposits marine deposits prevail. Delta progression became rapid during the high sea level period (6,000–5,000 bp) and was influenced by fluctuations in river sediment supply, neotectonics with areas of uplift and depression, relative sea level changes, and mangrove fringes, which promote sediment accumulation. During the past 4,000 years the delta front has prograded seaward by more than 90 km (more than 20 m/year) in the area along the Bassac channel, decreasing to about 50 km during the same period (about 12 km/year) to the northeast (with an overall average of 17–18 m/year between 5,300 and 3,000 bp and 13–14 m/year after 3,500 bp). The sediment discharge of the river during the past 3,000 years has been, on average, 144 ±36 × 106 tons/year, about the same as the present sediment discharge (Thi Kim Oanh Ta et al. 2002). There is no indication of any increase in sediment discharge caused by human activities (Van Lap Nguyen et al. 2000). At present much of the river sediment is supplied during periods of high river flow, and much sediment is flushed out into the coastal sea. During periods of low river flow turbidity reaches a maximum in the estuary, where much of the supplied sediment is ­deposited. The sand in the relict beach ridges and dunes dates mainly from the last interglacial period, and has been reworked several times by wind action. This is probably related to destruction of vegetation during climatic changes in the late Pleistocene (Murray-Wallace et al. 2002). The sediment brought into the coastal sea by the Mekong is largely moved southward under influence of

the northeast monsoon and goes around Mum Bar Bung into the Gulf of Thailand. South of the river mouths the coast down to Mum Bar Bung and the west coast from Mum Bar Bung up to Hon Chon consists of mangrove swamps and mudflats, with a short sandy beach west of Ca Mau. Hon Chon is composed of Palaeozoic limestone, and the offshore islands up to the The Chats archipelago consist of Palaeozoic and Mesozoic formations, with steep cliffs forming the coast. The large island of Phu Quoc has sandy beaches between promontories with rocky cliffs.

3.  The East Coast From Mui Vung ‘Mu (Cape St. Jacques) at the northeastern limit of the Mekong delta the coast north to Qui Nhon rises to mountains that reach a height of more than 2,000 m. Between Mui Dinh and Qui Nhon the coastline probably follows a N–S fault zone. The shelf is only 40–70 km wide, much narrower than along the other parts of the Vietnamese coast. Between Qui Nhon and Quang Ngai the Annamite Mountains come close to the coast, so that rocky spurs separate small alluvial plains and bays with short beach barriers or spits with dunes. The coast is indented, and there are many bays, as near Qui Nhon, Song Cau, To Bong, Ninh Hoa, NhaTrang, Cam Ranh, and Phan Rang. Spits face east-northeast to northeast, which is the direction of the strongest swell and the strongest monsoon winds. The capes are of granite, gneiss, and basalt, with rhyolite, dacite, and micaschists locally. From Vinh Quang Ngai north to Vinh there is a coastal plain, 30–50 km wide, regularly interrupted by mountain spurs reaching the sea and forming capes. Between the capes, which consist of granite or rhyolite (Mui Ron Ma, Mui Chon May Dong, Tien Shan Peninsula near Da Nang) or basalt (Mui Lay, Mui Batangan), there are broad, sandy barriers covered with low dunes and cut by small rivers flowing out into the South China Sea. Behind the beach barriers there is flat alluvial land, lagoons as in the Hue district and, landward of the barriers and lagoons, older beach ridges covered with dunes.

4.  The Northeast Coast From Vinh (18° 44' N) the coast turns northward, and consists of promontories of Cretaceous and Tertiary sandstone, siltstone, and conglomerate where NW–SE ridges reach the sea, separated by the alluvial plains of small rivers with estuaries bordered by beach ridges,

Vietnam

19.6

⊡⊡ Fig. 19.6.1 Beach ridges on the Mekong Delta. (Courtesy Geostudies.)

spits, and barriers. There are high, steep-sided islands of limestone, showing karstic shore weathering in Along Bay. The Song Hong (Red River) delta is low (less than 2 m above sea level), flat, and densely inhabited. Its coast is bordered by beach ridges, extensive mudflats, and mangrove swamps. The outflow of the river goes mainly through channels to the southwest, while the northeastern part of the delta is dominated by brackish tidal channels and mangrove swamps. Complexes of beach ridges and mudflats up to 30 km wide dominate in the centre and the southwest. The beach ridge crests are about 1 m above sea level, with low dunes, and are separated by depressions several hundred metres wide. The beach ridges are covered with vegetation soon after their for­ mation, and subsequently cultivated. They usually offer

sufficient protection for the delta farmlands against the high tides, whereas the mudflats are flooded during high tides. For further protection, dykes have been constructed since the eleventh century. They were primarily built for protection against flooding from the river and its branches, but also extend along most of the coast. The delta is prograding seaward by about 100 m/year in the south, the rate decreasing towards the northeast. North of the Song Wong river mouth there is hardly any progradation, most parts of the coast being more or less stable, with local erosion. There are indications that before the nineteenth century the rate of progradation in the south was slower, of the order of 30–40 m/year. North of the delta the coast consists of cliffs and rocky shores in Devonian schists, sandstone, and limestone, Permian limestones and Mesozoic sandstones, mudstones,

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and marls facing the Gulf of Tonkin to the southeast. An undulating Pleistocene landscape was submerged during the Holocene marine transgression to form numerous small islands and an indented coastline with inlets that have been partly filled with sediment and mangrove swamps. The Gulf of Tonkin is shallow, reaching little more than 50 m in depth at its centre, the continental shelf edge (200 m depth line) passing east of Hainan Island. There are indications of drowned valleys on the sea floor north of the Song Hong delta. An intricate coastline, fringed by many hilly islands, continues to the Chinese border at Mong Cai.

References Murray-Wallace CV, Jones BG, Tran Nghi, Price DM, Vu Van Vinh, Trinh Nguyen Tinh and Nanson GC (2002) Thermo-luminescence ages for a reworked coastal barrier, Southern Vietnam: a Preliminary Report. J Asian Earth Sci 20:535–548 Thi Kim Oanh Ta, Van Lap Nguyen, Masaaki Tateishi, Iwao Kobayashi, Susumu, Tanabe, Yoshiki, Saito (2002) Holocene delta evolution and sediment discharge of the Mekong River, Southern Vietnam. Quat Sci Rev 21:1807–1819 Van Lap, Nguyen, Thi Kim Oanh Ta, Masaaki Tateishi (2000) Late Holocene  depositional environments and coastal evolution of the  Mekong River Delta, Southern Vietnam. J Asian Earth Sci 18:427–439

19.7 Philippines

Eric Bird

1.  Introduction The Philippines archipelago consists of about 7,000 islands with a total area of about 300,000 sq km and an intricate coastline more than 13,000 km long. Extending from latitude 5° N to about 21° N, the islands are generally hilly or mountainous, with steep, locally-cliffed coasts and sectors of alluvial coastal lowland. The climate is tropical, ranging from humid (especially on the eastern side and in Mindanao to the south) to subhumid. Manila, the capital on the west coast of the largest island, Luzon, has a mean monthly temperature of 25°C in January rising to 27.8°C in July. In addition to the east to northeast trade winds (stronger in the north), there is a northeast monsoon (October–April) and a southwest monsoon (June– September). Ocean swell arrives from the Pacific Ocean, and to a lesser extent from the South China Sea, but within the archipelago wave action is generally weak, produced by local winds across often short fetches. Storm surges are found during occasional typhoons in summer and autumn. They move in from the Mariana Islands to the east and bring high winds, strong waves, and storm surges, mainly to the northern Philippines. Mean spring tide ranges in the Philippines are small: 0.8 m at Balabak, 1.1 m at Pankol, 1.4 m in Cebu, 1.0 m at Manila, 0.9 m in Mindanao, and 1.8 m at Davao on the south coast of Mindanao. From Luzon, in the north, the archipelago diverges southward on either side of the deep enclosed Sulu Sea. The Pacific coastline is bordered by the Philippine Trench, locally more than 10,000 m deep. Coral formations are extensive, both on the island coasts and as outlying reefs. There are many volcanoes. Some are active (Hibok, on Camiguin Island, north of Mindanao, erupted in 1951; Mayon in southeastern Luzon erupted in 1900, 1928, 1943, 1947, and 1968, and Mount Pinatubo, which had long been dormant, erupted in 1991); others, like Apo in southeastern Mindanao and Maricaban in Verde Island Passage south of Manila, are quiescent. Sands and gravels derived from the products of vulcanicity produce dark grey or black beaches, while coral formations yield beaches of white sand and gravel.

Neotectonic activity has been widespread, as evidenced by uplifted and deformed coral terraces (up to 360 m above sea level in Mindanao). Landslides have been triggered by earthquakes or very wet weather on coastal slopes in soft, deeply weathered materials and tsunamis have overwashed beaches and low-lying coastal areas. Coral reefs are widely distributed, mainly as fringing reefs along steep coasts. Mangroves are extensive in sheltered areas, but they have been reduced to less than 100,000 ha by clearance, mainly to establish brackish-water fishponds, which cover 176,000 ha (Juliano et al. 1982). In many areas deforestation has led to erosion on steep slopes and augmented siltation in estuaries and deltas.

2.  The Philippines coastline 2.1.  Luzon The north coast of Luzon, facing the Babuyan Channel, consists of a depositional plain narrowing east and west from the mouth of the Cagayan, the longest river in the Philippines (about 330 km). The wide estuary of the Cagayan at Aparri is fronted by shoals, but in general northeasterly waves (winter monsoon) have built long sandy beaches backed by dunes up to 8 m high, then corridors of swamp land, formerly lagoons. Behind these, are sectors of an older, inner sandy barrier. To the south, is hilly country, which runs out to Cape Engano and the steep-sided island of Palaui. To the north, the outlying Babuyan and Batan islands are mainly steep-sided volcanic peaks (e.g., Didicas). The Pacific coast is also steep and indented; the northern part, along the tectonically active Sierra Madre, receiving strong ocean swell, which has built and shaped curving bay beaches. To the southeast, Luzon has a more indented coastline, with many sheltered gulfs, some edged with mangroves. In Larap Bayon, the south coast sedimentation has been augmented by the inflow of silt and clay washed down from iron ore mines in the nearby hills (Rabanal and Datingalang 1973). Near Legaspi the beaches are black sand of volcanic origin, mainly from the Mayon

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volcano, but outlying islands are often reef-fringed, with white sandy beaches. Verde Island Passage is a narrow strait bordered by steep coasts, with the island volcano of Maricaban in the middle. A low-lying isthmus separates the muddy lake of Laguna de Bay (which occasionally receives brackish inflow in dry summers) from Manila Bay. South of Manila the Cavite spit has grown northward, sheltering Bacoor Bay, permitting the growth of deltas, the building of fishponds and land reclamation. Manila Bay occupies a sheltered embayment with a large delta built by the Pasag, Pampanga, and Bulacan rivers growing from its northern shores. These have wide bordering mudflats, and the coastal fringe has been intensively used for fishponds and land reclamation. To the north, the volcanic Bataan Peninsula has coves at the mouths of radial valleys and intervening cliffy sectors. North to Cape Bolina, the spectacular Zambales coast is steep, rugged, and reef-fringed, with small plains mainly at river mouths. Inland is the volcano, Mount Pinatubo. The west coast of Lingayen Gulf (>Fig. 19.7.1), bordering the highlands, is steep with rocky cliffs fringed by coral reefs, some of which are emerged and dissected, as in the Hundred Islands (>Figs. 19.7.2 and >19.7.3). To the south of Nanatian is a gravelly beach (>Fig. 19.7.4) with inputs of pebbles and cobbles from streams that drain the uplands. It becomes sandy westward along the south coast, which consists of a long curving sandy beach (>Fig. 19.7.5), prograded by the onshore movement of sand bars from a shallow sea floor The beach is interrupted by the mouths of rivers (Calmay, Beud, Dagupan) and backed by numerous low sandy beach ridges, then swamplands largely in use for fishponds or rice cultivation. The east coast of Lingayen Gulf has small lobate deltas built by the Naguilian and Samara Rivers which have outlines indicating southward longshore drifting of sand lobes, culminating in a large spit south of Santa Barbara. San Fernando has a sandy tombolo linking a cliff-edged island, and to the north, at Luna, a historic coastal watchtower has being undermined by wave scour. The west coast of northern Luzon is generally steep, with fringing coral reefs and some beaches. Emerged coral sectors have typical notch and visor profiles.

Galera. Many short, steep streams flow down forested slopes to rocky shores. Much of the west coast is beachfringed and consists of swampy plains: a long, gently curved beach runs from Talabasi Point south to the Sablayan tombolo, breached only by the mouths of small rivers. Farther south, the Mongpong River has built a substantial lobate delta, and the Lumintau River opens to the sea by way of a wide estuary. Sand washed down the braided Bagsanga River drifts mainly southward as beaches, culminating in the long Caminawit spit in front of mangrove-fringed Mangarin Bay. In the southwest the Caguray River has built a large swampy delta, constricting the strait in the lee of the steep high island of Ilin. The south coast of Mindoro has several hilly promontories, where submergence of a much-dissected volcanic topography has produced numerous coves, inlets, and outlying islands east of Bulalacao Bay. Fringing coral reefs and white sandy beaches are present locally here. Much of the east coast is low-lying, a fluviodeltaic plain with fringing beaches of grey sand. In some sectors, multiple beach ridges are separated by swampy swales, and spits flank the mouths of small rivers. Low-lying Duyagan Point is the apex of a large delta built by the Saquisi River and its distributaries. To the north, spiky volcanic hills form the Dumali Peninsula alongside Pola Bay, where a sandy beach is backed by extensive mangrove swamps. The northeast coast is generally beach-fringed, with some coral reef sectors. Marinduque Island, to the east, is a reef-edged volcanic upland with mainly white sandy beaches.

2.2.  Mindoro

Panay Island has many features similar to Mindoro. Its west coast is largely beach-fringed with some surrounding reef sectors fronting the steep cordillera ranges, and deep water close inshore. Several rivers deliver sand to the coast, where prograded beach-ridge plains have been formed. Southward drifting from the Sibalom River has built the foreland of Tubigan Point at San Jose. In the

Mindoro is a hilly to mountainous island with steep forested coasts in the north, facing the Verde Island Passage, and northwest from Cape Calavite down to Mambuarao. In the north, the hilly peninsula behind Escarceo Point has an irregular outline, sheltering the reef encircled Port

2.3.  Palawan Palawan, an elongated high island southwest of Mindoro, is flanked by submerged barrier reefs, and has steep coasts with many white sandy beaches. Near El Nido, there are high cliffs of dark limestone, but volcanic formations predominate, and fringing reefs border headlands along the embayed coastlines.

2.4.  Panay

Philippines

⊡⊡ Fig. 19.7.1 Lingayen Gulf. (Courtesy Geostudies.)

cliffs

SAN FERNANDO

severe erosion

L I N G A Y E N

R

Bauang

seasonal sand barrier at mouth

n Naguilia

sand lobes moving southward R

ba Ca

wide sandy beaches

Modern coral reefs

G U L F

Samara

erosion (sea wall)

R

Santa Barbara Agoo

groynes

southward drift to growing spit Damortis HUNDRED ISLANDS emerged and dissected coral leaf

submerged sand spit low cliffs cut in emerged coral reef San Fabian

steep, rocky coast Nanatian

gravelly streams from uplands

ow all sh d gravel beach in beh ches a e b Sandy

LINGAYEN

a se

DAGUPAN

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⊡⊡ Fig. 19.7.2 The Hundred Islands, emerged coral reefs in Lingayen Gulf. (Courtesy Geostudies.)

⊡⊡ Fig. 19.7.3 Notch and visor profiles on limestone stacks in the Hundred Islands, Lingayen Gulf. (Courtesy Geostudies.)

southwest, reefs flank rugged promontories and outlying islands, but the south coast is again sandy, with an eastward drift towards the port of Iloilo in the strait behind the hilly island of Guimaras. Southeastern Panay is essentially a composite deltaic plain built by the Iloilo and Jalauod ­rivers, but the east coast is hilly and embayed,

with many reef-fringed headlands and islands, some of which are uplifted coral reefs. The spring tide range here is slightly higher than in the Philippines, generally attaining 2.4 m in Concepcion Bay. As a result, tidal mudflats are extensive, and fishponds have been built on these, and in the ­mangrove fringe. In the north the Panay River has

Philippines

19.7

⊡⊡ Fig. 19.7.4 Gravel beach south of Nanatian, Lingayen Gulf. (Courtesy Geostudies.)

⊡⊡ Fig. 19.7.5 Prograded beach of dark volcanic sand, south coast of Lingayen Gulf. (Courtesy Geostudies.)

built a large delta, the seaward edge of which has sandy beach ridges, backed by broad mangrove and nipa swamps intersected by distributaries and tidal channels. Long sandy beaches and beach-ridge plains extend west from here, forming paired spits at the mouth of the mangrove-fringed Batan Lagoon and curving out to the Aklan Delta.

2.5.  Negros Negros Island is mountainous and forested, culminating in the Canlaon volcano (2,450 m) with intensively farmed coastal plains in the north and west, where extensive reefs and shoals front mangrove swamps and beach sectors. Low-wave energy conditions predominate along the shores

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of Guimaras Strait, where the Malogo and Bago rivers have built substantial deltas, with low sandy ridges and muddy nearshore zones. To the south, the coast becomes hilly and more exposed past Sojoton Point, and reefs fringe  generally rocky shores: there is a coral lagoon at Tinagong Dugat, near Sipalay. The east coast, bordering Tanon Strait, is also generally steep, with extensive reefs often backed by mangroves, as at Bais. There are many uplifted reef sectors. Outlying Siquijor Island to the southeast is high, hilly, and surrounded by reefs and beaches of coarse coralline sand.

2.6.  Cebu The coastlines of Cebu, an elongated high island with many raised coral reefs, and Bohol, less mountainous, show features similar to those of eastern Negros, with generally low wave energy on the shores of narrow straits. Karstic coastal features are found on emerged coral sectors. Where wave action is stronger, as in northeastern Cebu, there are longer and bolder beaches, and seaside resorts have developed such as at Liloan Beach.

2.7.  Masbate, Samar and Leyte Masbate, to the north, is relatively low-lying and has an embayed coastline, with inlets such as Nin Bay indicative of recent submergence. Samar and Leyte to the east, continue the forms of southeastern Luzon, with their held and often cliffed eastern coastlines being subjected to easterly ocean swell and to recurrent typhoons and surges. The western shores, on the other hand, are more sheltered, with reef-fringed promontories and islands and mangrove-edged muddy bays and inlets. Although segments of curved bay beaches are found on the ocean coasts, there are many cliffed promontories, reef areas, and outlying rocky islands. Emerged reef terraces are found near Guiuan in

southeastern Samar. Leyte Gulf is penetrated by ocean swell that breaks on long sandy beaches on the east coast of Leyte (>Fig. 19.7.3). Northward drifting here has built a large recurved spit at Cataisan Point, near Tacloban.

2.8.  Mindanao Mindanao, southernmost of the large islands, is mountainous. Off its north coast is the active Hibok volcano on Camiguin Island, which has deposited lava and ash on the coast during successive eruptions, notably in 1951. Mambajao, the town on its northern shore, is built on an earlier lava flow. Lava boulders litter the shore, interspersed with beaches of black volcanic and white coralline sand. On Mindanao, beach ridges have been prograded at the mouths of rivers such as the Agusan, entering Butuan Bay, and the streams flowing into Macajalar Bay at Cagayan de Oro. Mangroves back sheltered bays and inlets, and form swamps near river mouths. In the southwest, the Zamboangan Peninsula has steep, reef-fringed coasts and beaches of shelly and coralline sand, and similar features are seen on Basilan and the chain of high islands running southwest towards Sabah. In the southeast, there are beaches of black volcanic sand on the shores of Davao Gulf, beneath the slopes of Mount Apo. Landslides are common on steep coastal slopes here, and there are high cliffs on Cape San Agustin.

References Juliano RO, Anderson J, Librero AR (1982) Philippines: perceptions, human settlements and resource use in the coastal zone. In: Soysa C, Lin Sien Chia, Collier WL (eds) Man, land and sea. Agricultural Development Council, Bangkok, pp 219–240 Rabanal HR, Datingslang BY (1973) Problems of pollution and siltation from mines in some Philippine waters. In: Costin AB, Groves RH (eds) Nature conservation in the Pacific. Australian National University Press, Canberra, pp 263–270

19.8  Indonesia    Otto Ongkosongo   

1.  Introduction

Much of Indonesia has a humid tropical climate, with high temperatures and rainfall, but there are variations related to the position of the Intertropical Convergence Zone of unstable air and heavy rainfall, which migrates north and south over Indonesia, crossing the Equator in May and November each year and reaching about 15° S in January. When this zone is to the south, there are prevailing westerly winds and heavier rainfall, although northeasterly trade winds reach some northern coasts; when it moves north, southeasterly winds bring drier conditions, especially along the southern coasts. Winds are generally light to moderate. Wave action in Indonesian waters is largely generated by local winds, gentle in the equatorial zone but stronger on the northern and southern coasts subject to north­­ east and southeast trade winds, respectively. Ocean swell moves in to the southern coast from the Indian Ocean and to the northern coast from the southwest Pacific. In general, wave energy is low. Tropical cyclones do not reach

Indonesia consists of about 18,000 islands, with an intricate coastline of just over 80,000 km (>Fig.  19.8.1). In terms of global tectonics, the Indonesian archipelago occupies the collision zone between the Indo-Australian, Pacific and Eurasian plates and is a region of continuing instability marked by frequent earthquakes and volcanic eruptions.It includes mountainous areas of Tertiary and Quaternary uplift, augmented by large volcanoes, and in coastal reg­ions, there is widespread evidence of uplift and depression, often accompanied by tilting and faulting, in Pleistocene and Holocene times. Coral reefs are numerous and extensive in Indonesian waters, and in many places, they have been raised out of the sea by tectonic movements as in Sumatra, Java and Irian Jaya (West Papua). The southern part of Indonesia is controlled by an active subduction zone which is marked by deep trenches off Sumatra and Java. ⊡⊡ Fig. 19.8.1 Indonesia: Location Map. (Courtesy Geostudies.) Kota Bharu

NIAS

Str Sandakan

Kota Kinabalu

CAROLINE

EAST MALAYSIA

Celebes Sea

Tarakan

SARAWAK Kuching

PACIFIC

Manado KALIMANTAN

r Ka

sar

ra it

S

kas

ND

St

LA

IS

BANGKA Palembang

a

nd

Su

BELITUNG Java Sea it a I r St Djakarta

Bandjarmasin

N

Bandung

D

O

CELEBES

N

E

Udjung Pandang

S

Cocos Is. (Australia)

OCEAN

EQUATOR

Jayapura

SULA

I

Wewak

A

BURU

PAPUA NEW GUINEA

IRIAN JAYA

CERAM

BUTUNG

MADURA Surakarta Flores Sea Surabaja INSET MAP J A V A BALI LOMBOK Jogjakarta g n FLORES la Ma Denpasar SUMBAWA Christmas I. Kupang SUMBA (Australia)

N

HALMAHERA Molucca Sea MOLUCCAS

Balikpapan

a

AI

at

W

im

TA

EN

M

S U MA T R A Padang

ISLANDS

SABAH

BRUNEI

it

Medan

bac

Ma

SIMEULUE

Bala

MALAYA Ipoh WEST MALAYSIA St Kuala ra it of Lumpur M al ac ca SINGAPORE

Stra

Penang

Banda Sea WETAR ALOR

RU ISLANDS JAMDENA

Dili TIMOR

Torres Strait

Arafura Sea

INDIAN Timor Sea

OCEAN 0

500

1000 km

Weipa

Darwin

AUSTRALIA

Gulf of Carpentaria

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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platforms and coral reefs. These are seen along the southern shores of Sumatra and Java, in western and northern Sulawesi (including Gorontalo Province), on the south coasts of the eastern islands and in the north of Irian Jaya. Limestone cliffs, in particular, show the effects of intense physical, chemical and biological weathering, with notches and caves. Beaches are extensive in Indonesia, some derived from fluvial sands and gravels, others from cliff erosion and still others from the erosion of calcareous material from fringing coral reefs. Beach sediment derived from volcanic rocks are typically black or grey; those of coralline origin are white or yellow. In the granitic zone of the Riau, Bangka and Belitung Islands, white quartz sands dominate ­beaches. Sandy backshores are colonised by coastal vegetation, not­ably Ipomoea pes-caprae and Spinifex littoreus, then coconut and casuarina trees. Coastal dunes are poorly developed in the humid tropics, but on the southern shores of Java and Sumatra, prograded beaches are backed by dunes, some of which carry woodland vegetation. Some parts of the Indonesian coast have prograded by the deposition of lava and ash from volcanic eruptions, as in the Krakatau Islands. Large quantities of pyroclastic sediment have been transported down to the coast from active volcanoes, such as Merapi in southern Java or Agung in Bali, and from the erosion of dormant or extinct volcanoes as on Manado Tua Island north of Menado.

Indonesia, but waves generated by such disturbances are occasionally transmitted into Indonesian waters, especially along the southern coasts. Tidal movements result from impulses that arrive from the Pacific Ocean by way of the South China Sea and the Philippine Sea and from the Indian Ocean through the Straits of Molucca and along the southern coasts to the Timor and Arafura seas. Mean maximum tide ranges (>Fig. 19.8.2) are generally less than 2 m and only exceed 6 m locally on the southwest coast of Irian Jaya. Tides are complicated by wind action, especially in the trade wind zones, and by the effects of tectonic and volcanic disturbances that generate tsunamis. The explosive eruption of Krakatau in Sunda Strait in 1883 generated a tsunami up to 30 m high on adjacent coasts and lesser surges all around the Indonesian coastline. Tectonic tsunamis sometimes res-hape the coasts in the subduction zones and may also near faults in the internal waters of Indonesia, as on Maumere and the surrounding islands, notably Babi Island. The Boxing Day 2004 earthquake south of Sumatra generated a large tsunami that caused extensive damage and sea flooding, particularly in Aceh province, and another earthquake in March 2005 caused uplift and subsidence on Nias and nearby islands south of Sumatra. High parts of the coast are generally steep and forested rather than cliffed, but where there is relatively strong wave action (including swell) from the Indian Ocean or the Philippine Sea, there are bold cliffs, fronted by shore

⊡⊡ Fig. 19.8.2 Variations in tide range around the Indonesian archipelago. (Courtesy Geostudies.) 1.2 0.4 0.4

0

1.4

500

1000km

N

1.9 2.0

2.2

U

1.7

M

1.1

T R A

1.9 2.2

IN

DO

1.0 0.9

NE

KALIMANTAN

0.8

0.5 0.6 1.1

SI

1.5

1.5 1.5 0.8

1.0

2.8

2.2

0.9 0.7

A 1.0

CELEBES SEA

2.5

3.8

S

0.4 0.6 0.6 0.8

JAVA 0.4 0.4 JAVA 1.5

AN

2.4 2.1 1.0 1.4 SEA 0.8 0.6 0.5 1.7

1.1

1.5 1.0 0.7

MEAN SPRING TIDE RANGE (metres)

OC

EAN

0.9

1.8 1.0

0.8 0.9

1.2

0.6

1.3 1.4

0.7

1.5

0.6

0.9

0.8

1.0

1.4

BANDA 1.4 1.2 FLORES SEA 1.5 1.0 1.8 1.2 O R 2.0 2.4 2.6 TIM 2.1 TIMOR

SEA 6.4

10.5

1.7

1.4

1.5

1.9 2.6 4.2

1.6

5.4 3.4

ARAFURA

SEA

2.8 5.5

5.7 6.6

0.7

IRIAN JAYA

1.0 SEA

0.9

1.0

1.2 1.6

MOLUCCAS

0.9 SULAWESI0.4

1.7 0.6 0.7 0.5 0.5

0.6

1.7

2.0

2.0 1.9

1.7 AUSTRALIA 1.1

Indonesia

Apart from volcanic sources, rivers draining mountainous uplands of conglomerate, sandstone and shale have carried gravel, sand, silt and clay to the coast to form deltas and coastal plains as in East Sumatra, North Java, Kalimantan and Irian Jaya. Prograded coasts are commonly marked by multiple beach ridges alternating with swales, as on the central south coast of Java. Many features of coastal topography are related to the effects of deep weathering under humid tropical conditions: there are steep vegetated coastal slopes with recurrent slumping on deeply weathered rock outcrops, with cliffs confined to occasional bedrock headlands, and extensive deltas have been built by rivers that bring down an abundance of fine-grained sediment from sloping hinterlands, as in eastern Sumatra, northern Java, Kalimantan and the south of Irian Jaya. At least 50 mangrove species are found in Indonesia, and mangrove-fringed coasts are extensive, especially in the northeast of Sumatra, alongside the estuaries of southern Kalimantan, and in southern Irian Jaya. Mangrove swamps develop on mud rich in organic matter, which darkens the sediment. In densely populated areas, as in northern Java, the mangrove fringe has been largely cleared to make way for brackish-water fishpond (tambak) construction. Coastal features have also been modified by human activities where the sediment yield from rivers has increased as the result of accelerated erosion in their deforested and cultivated catchments. Some parts of Indonesia, however, have a semi-arid climate that results in thin soils and exposed rock outcrops, as in Nusa Tenggara Timur and Nusa Tenggara Barat. Coral reefs are widespread, except off the mouths of rivers where the sea is freshened and made turbid by water and sediment discharge and where active volcanic activity has inhibited coral establishment. There are many reeffringed islands and reefs that enclose lagoons with central islands, and true atolls, especially in the Flores and Banda Seas. The major barrier reefs and atoll reefs are the Great Sunda Reef, rising from the submerged shelf margin southeast of Kalimantan, the Takabonerate reef south of Sulawesi and the reefs that curve out toward the islands of Batu and Banyak off Sumatra. Algal rims are better developed on reefs along the oceanic southern shores than in the inner seas. Many small islands are formed by coral reefs such as the Seribu Islands north of Jakarta or as reeffringed islands such as Karimunjawa Islands north of Semarang. The record of changing sea levels in Indonesia has been complicated by uplift and subsidence of the land by Quaternary tectonic movements. There are terraces and former coastlines at various levels, and Tjia (1975) was of

19. 8

the opinion that the higher Holocene stillstand, well ­documented on the east coast of peninsular Malaysia, could also be traced through much of Indonesia. During the Last Glacial low sea level phase, Java, Sumatra and Kalimantan occupied an enlarged Malaysian peninsula, separated by deep straits from Sulawesi and the eastern islands, Australia then being linked to New Guinea. On Timor and Atauro, stairways of emerged coral terraces have been dated and related to land uplift at rates of up to 0.5 m per thousand years during the Late Quaternary sea level oscillations, and there are emerged reefs on Sumbawa Islands.

2.  The Coasts of Indonesia 2.1.  Sumatra The western and southern coasts of Sumatra are mostly steep and cliffed, with intervening lowlands dominated by beach ridge plains, as at Tapak Tuan, Sibolga, Padang (Verstappen 1973), Padang Bai (Bengkulu) and Teluk Lampung. There are many pocket beaches that were formed and nourished with sediment from eroding cliffs, rivers and fringing coral reefs, with some local shell contributions. Rivers are generally short compared with those flowing to the east (Malacca Strait) coast, and some descend over waterfalls on or inland from coastal cliffs. An anticlinorial mountain chain with associated volcanoes runs NW–SE through the island, dissected by the active Semangko fault system. It is bordered to the south of Sunda Strait by the Java trench and other roughly parallel submarine fault systems between the mainland of Sumatra and the island arc of Sinabang, Nias and Mentawai, which have most of the cliff and bluffs along the southwest coast. There was local uplift and depression on these islands, as well as tsunami submergence, during the March 2005 earthquake. The rivers that flow northeast and east coasts to the shallow Straits of Malacca have formed a broad depositional lowland with extensive wetlands bearing fresh water forest in the upper coastal plain and mangrove swamps in the lower coastal plain. Beach ridges have been found up to 150 km inland, and the coastline of the Jambi area has prograded by up to 75 km in the past century, narrowing to the southeastern tip of Sumatra to form lineaments. In Lampung Bay, mining of a coastal hill has exposed the inner part of a volcanic neck as an artificial bluff. The Peusangan Delta in northern Sumatra shows stages in growth and decay related to river capture. The

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beheading of Djuli River by river capture was followed by erosion of its delta, while the pirate stream, the Peusangan River, built a new delta to the east in two stages. Beach ridges enable the sequence of delta growth and decay to be traced (>Fig. 19.8.3). When the Batang Arau river in Padang was diverted to the south, the former delta to the north was eroded, and the coastline had to be maintained by building revetments and sea walls. Erosion has also occurred on the Tapak Tuan coastal plain in Aceh Province. The Boxing Day 2004 tsunami caused extensive sea flooding and structural damage in Aceh Province. Farther south rivers such as the Rokan, Kampar, Indragir and Musi-Banyuasin open to broad tidal estuaries fringed with mangroves and occupied by low-lying islands. Ports which were close to the sea in the fifteenth century, such as Palembang and the more recent Bagansiapi-api, are now far upstream, indicating that there has been rapid coastal progradation. However, rapid clearing and exploitation of the mangroves have led to erosion on many parts of the coast, as at Musi-Banyuasin, and deforestation has caused severe shore erosion in East Lampung.

Growth of a longshore spit has partly enclosed a coastal lagoon in Padang Bay. Sediment from an inflowing river has been deposited in the lagoon, and there has been shoaling of the lagoon mouth, narrowed by the growth of the spit. The lagoon had been used as a port, and a long jetty was built to maintain the lagoon mouth, but this has led to updrift accretion and downdrift erosion. Emerged fringing reefs up to 3 m above sea level are found on the Aceh coast and nearby islands south of Lampung Bay, and the shores of Semangka and Lampung Bays, and along the coast of Sunda Strait, are strewn with wave-deposited coral boulders. Parts of the mangrove coast are still prograding, but some have been retreated, even on relatively protected coasts as in the Dumai area. In the large Rokan, Kampar and Musi-Banyuasin estuaries, mangroves occupy low depositional islands, especially where there is protection by offshore hilly islands. Within mangrove swamps, there are crab mounds that rise several centimetres above the mud surface. Many mangrove swamps have been converted into intensive fish or shrimp ponds (tambak) that no longer receive river-borne silt because the enclosing banks exclude flood waters.

⊡⊡ Fig. 19.8.3 Changes near the mouth of the Peu­sangan River, northern Sumatra, following a river capture. Beach ridge patterns indicate the trend of an old delta north of Bireuen and two stages in the development of a new delta to the east: At A, a lobe that has been truncated by erosion. At B, a developing modern delta. (Courtesy Geostudies.)

19. 8

Indonesia

2.2.  Krakatau

Rias are common in the Riau Islands, produced by the Late Quaternary marine transgression invading the lower parts of river valleys. They are generally mangrove fringed. On Batam Island, Duriangkang Bay is one of several bays that have been enclosed by artificial dykes to form fresh water reservoirs. There has been tin mining in granitic areas on some of the Riau islands, together with Singkep, Bangka and Belitung Islands, and quartzose waste has formed beaches, spits and barriers and a small artificial island east of Bangka. There has also been sand quarrying on these islands for export to Singapore and Malaysia as ­reclamation material. As a result, some of the small islands have disappeared, others have been intensively eroded, and many fringing coral reefs have been destroyed. The granitic coasts are dom­­inated by quartz sand beaches, while coasts on metasedimentary rocks have rocky cliffs and muddy shores. The island of Karimun, west of Singapore, has granite-weathering features, including tors and rillenkarren, rising from deeply weathered mantles on coastal slopes.

⊡⊡ Fig. 19.8.4 The Krakatau Islands. (Courtesy Geostudies.)

Krakatau, an island volcano in Sunda Strait, erupted explosively in 1883, leaving residual steep-sided islands (>Fig. 19.8.4), Panjang, Sertung and Rakata (>Fig. 19.8.5). Subsequently, a new volcano, Anak Krakatau, has formed within the caldera (>Fig. 19.8.6). Sunda Strait is bordered by steep volcanic coasts, and the effects of the 1883 eruption of Krakatau (Symons 1888) are still evident. The eruption left a caldera of irregular outline, 7 km in diametre and up to 250 m deep, which has been modified by rapid recession of cliffs cut into soft pumice on the residual islands, which are now forested. The 1883 tsunami destroyed reefs and mangrove swamps, and the large coral boulders that were thrown up on the shore have since been undercut by up to 35 cm by shore weathering processes. By 1927 the new volcanic island, Anak Krakatau, had formed in the caldera, and was expanding in area and growing in height. It reached a height of 198 m by 1983, and in 2000 it was about 230 m high. It is dissected by deep gravelly gullies, and has coasts

ng

accre

ti

e ro d in

g

spit

Panjang

cliff

f lif

af lav

c cli

ff

s low

.

0m

20

Anak Krakatau

cliff

sandy cusp

cliff

Sertung

sandy cusp

cliff

Karang Serang

.

250 m

Rakata

cli 0

1

2

3

4

5 km.

ff

ff

caldera

cli

wa ll

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⊡⊡ Fig. 19.8.5 The steep cliff on Rakata is the wall of the crater formed by the 1883 explosion. (Courtesy Geostudies.)

⊡⊡ Fig. 19.8.6 Anak Krakatau, the new volcano developing in the Krakatau caldera. (Courtesy Geostudies.)

with promontories of lava (>Fig. 19.8.7) and cliffs cut in volcanic ash (>Fig. 19.8.8). Some plants and animals have colonised the new island.

2.3.  Java Java is the most densely populated island of Indonesia. It extends about 1,000 km from west to east and is up to

200 km wide. The higher parts of the island are formed by a chain of active and dormant volcanoes that run along its length, and have supplied large quantities of sediment as a result of frequent eruptions (as from the Merapi volcano), or erosion of the volcanic mountains. The drainage divide follows the chain of volcanic peaks. Other sedimentary or volcanic mountains and hills also contribute large quantities of sediment, aided by intensive humid tropical weathering processes. Rainfall of up to 7,000 mm/year produces

Indonesia

19. 8

⊡⊡ Fig. 19.8.7 Lava promontory of Anak Krakatau. (Courtesy Geostudies.)

⊡⊡ Fig. 19.8.8 Cliffs cut into pyroclastic deposits on the south coast of Anak Krakatau. (Courtesy Geostudies.)

continuous flow in many rivers that have deposited sediment to form wide coastal plains and growing deltas. Rapid deforestation, especially in mountainous areas, has intensified erosion of the land surface and increased the yield of sediment to rivers, some of which show channel and river mouth shoaling. The west coast of Java, between Anyer and Labu-han, is mainly formed by an emerged fringing coral reef, the seaward margin of which is being eroded in many places. Coral boulders are widespread along the shore, swept in

from fringing coral reefs by the 1883 tsunami and deposited on the shore platform (>Fig. 19.8.9) and the coastal plain. There are erratic boulders of volcanic agglomerate and breccia in addition to the coral blocks. On the north coast of Java, deltaic plains have been built out into the relatively low wave-energy microtidal Java Sea by sand and silt-laden rivers. During the past century sectors around the mouths of rivers and canals have prograded, while sectors no longer receiving fluvial sediment have been subject to erosion. In several cases, river

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⊡⊡ Fig. 19.8.9 Coral reef blocks deposited on the shore at Anyer, Java, by the 1883 tsunami. (Courtesy Geostudies.)

eroding

0

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mouths have been diverted either naturally like the Cimanuk in 1947 or as the result of the cutting of canals, as on Ciujung and Ciwandan deltas, initiating new delta growth. The coastline of Jakarta Bay shows stages in delta progradation, especially on the eastern shores around the mouths of Citarum River distributaries. Earlier coastlines are traceable from patterns of beach ridges in the Jakarta region and also in the deltaic plains to the east. In re­­ cent  years, the Citarum Delta has grown northwestward (>Fig.  19.8.10) and erosion has become extensive on its northern flank, partly because river outlets have migrated westward and partly because sediment yield has ­diminished

N

⊡⊡ Fig. 19.8.10 The Citarum delta, northern Java, formerly built a delta northward, but this is eroding away following a swing to the northwest, where the modern delta is developing. Dam construction upstream has reduced the sediment yield here and slowed delta growth. (Courtesy Geostudies.)

following the completion of the Jatiluhur Dam upstream in 1970 (Bird and Ongkosongo 1980). Erosion has damaged shrimp ponds (tambak) that has been constructed near the delta margin. Northwest of Jakarta Bay are the coral reefs and cays of the Thousand Islands, many of which have changed in outline during the past century. Some have enlarged by the accretion of sand on cays and shingle on northeastern (windward) ramparts; others show erosion or lateral displacements related to variations in the local wind regime. Some are subsiding, others are rising and some have been eroded as the result of removal of shingle and coral.

Indonesia

To the east of Jakarta, the Citarum, Cipunegara and Cimanuk, with smaller intervening rivers, have built up a major confluent deltaic plain. The Cimanuk has built a new delta following its diversion northeastward during a flood in 1947. There has been subsequent growth along three distributaries, with further branching developing as the result of median shoal formation in the river mouths. The older delta to the west is eroding away, as is the eastern flank, which curves out to a low-lying promontory, Cape Ujung. This was thought to be a relic of an earlier subdelta, but there are no relics of former channels leading in this direction. Another suggestion has been that it is related to a buried or nearshore reef structure, but no ­evidence exists for this either. The promontory occurs where an earlier beach ridge system was truncated by the present coastline. On the smaller deltas to the east (Bangkaderas, Bosok, Pemali, Comal and Bodri), Hollerwöger (1964) traced the stages in historical evolution and noted an acceleration of growth after 1920, related to increased sediment yield due to clearance of forests on steep hinterlands and the introduction of farming to these areas. In detail, there have been sectors of advance and retreat related to the changing locations of river mouths. Near Jepara a new delta formed at the mouth of a canal cut in 1892 to divert the outflow from the Kedung River, and the Solo delta has grown rapidly. There has been substantial progradation of the coastal plains around Surabaya in recent centuries, narrowing Surabaya Strait, which shows tidal scour features indicative of strong current action. ⊡⊡ Fig. 19.8.11 Notched coastal slope at Baron, southern Java, behind a shore platform that has a veneer of black volcanic sand. (Courtesy Geostudies.)

19. 8

The south coast of Java has cliffed promontories and limestone cliffs, often fronted by shore platforms and fring­ ­ing reefs. The limestone cliffs have basal notches and visors behind the shore platforms (>Fig. 19.8.11), and as these are as well developed on the lee shores of stacks and islands that have been shaped largely by solution processes. Sand supplied by rivers has been deposited to form extensive beaches, with local cuspate forelands, as behind Nusa Barung Island and tombolos, as at Pangandaran. The wide coastal plain near Yogyakarta is related to an abundant fluvial sediment yield from the steep hinterland, including sand and gravel brought down by rivers from the Merapi volcano, particularly after eruptions, when lahars flow down slopes and river discharge is torrential. Dark grey sand is delivered to river mouths and spread along the shore. Formation of protruding deltas of the kind seen in northern Java is prevented by relatively strong wave action, the shoreline being almost straight past the mouths of several rivers, each of which shows westward deflection by longshore drifting produced by waves generated by southeasterly winds. To the west of Parangtritis, the seaward margin consists of beach ridges and active dunes up to 30 m high, driven inland by the southeasterly winds. These dunes were evidently mobilised when their scrub and woodland cover were destroyed by sheep and goat grazing and the harvesting of firewood. They are quite anomalous in a humid tropical environment. On the south coast of Central Java the westward growth of longshore spits due to the oblique arrival of waves generated by southeasterly winds may block river

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⊡⊡ Fig. 19.8.12 Beach ridges on the coastal plain east of Cilacap. (Courtesy Geostudies.) SOUTH SERAYU RANGE SOUTH SERAYU RANGE

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⊡⊡ Fig. 19.8.13 Segara Anakan, a mangrove-fringed lagoon. (Courtesy Geostudies.)

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19. 8

Indonesia

mouths for some weeks or months during the dry season, forming temporary coastal lagoons. The spits are usually breached when river outflow increases in the wet season. The coast west to Cilacap has prograded intermittently, with the addition of successive beach ridges which are interrupted by alluvial tracts where rivers such as the Serayu flow across them (>Fig.  19.8.12). At Cilacap the coastal plain extends behind a high limestone ridge, and extensive mangrove swamps, intersected by tidal channels and creeks, border the broad, shallow estuarine lagoon of Segara Anakan. Inflow of large quantities of silt, especially from the Citanduy River, is reducing the depth and extent of this lagoon system and promoting mangrove ­encroachment (>Fig.  19.8.13). Farther west, sectors of steep coast alternate with beach-ridge plains along and towards Java Head, at the southern entrance to Sunda Strait.

2.4.  Kalimantan Geomorphologists have given very little attention to the coasts of Kalimantan. Swampy plains are extensive, but rates of progradation have not been documented. Many river mouths are estuarine, especially on the east coast

2.5.  Sulawesi Sulawesi has generally steep coasts, with terraces (including emerged, tilted, and warped coral reefs) up to 600 m above sea level indicating tectonic uplift. Rivers are short and steep, with waterfalls and incised gorges and only minor depositional plains at their mouths. Volcanoes are active locally, notably on Menadotua off the north of the island.

2.6.  Madura The island of Madura is notable for its straight northern coast which may follow a fault line.

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⊡⊡ Fig. 19.8.14 Morphological features of the Mahakam delta on the east coast of Kalimantan. The high-tide shoreline is intricate and dynamic, but the delta really extends seaward as a submerged sedimentary lobe with a steep outer slope, here indicated by the 10- and 20-fathom lines. (Courtesy Geostudies.)

where the tide range is more than 2 m, but the Mahakam has built a major delta (>Fig. 19.8.14) formed largely of coarse, sandy sediment derived from fluvial erosion of sandstone ridges near Samarinda, with swampy areas be­­ tween the distributaries. On the west coast, only the Pawan and Kapuas rivers have carried sufficient ­sediment from their extensive catchments to build protruding deltas.

0°50 S

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⊡⊡ Fig. 19.8.15 The limestone cliff at Uluwatu on the southeast coast of Bali. (Courtesy Geostudies.)

⊡⊡ Fig. 19.8.16 Fringing reef on the east coast of Bali, backed by a beach of coralline sand derived from the reef. (Courtesy Geostudies.)

2.7.  Bali Bali consists partly of volcanic terrain, with active volcanoes such as Agung periodically generating lava and ash deposits that are washed down to the sea by rivers,

and partly of coralline limestone formations, which are cliffed on the Bukit Peninsula to the southeast (>Fig. 19.8.15). Beaches border much of the island, with coral sands behind fringing reefs (>Fig.  19.8.16) contrasting with dark grey sand derived from cliff outcrops

Indonesia

of volcanic ash, or carried down to the sea by rivers draining the volcanic terrain. The spit at Gilimanuk shows evidence of alternating growth and truncation, and in the Sanur area spits and barrier islands enclose a broad, mangrove-fringed tidal embayment north of the sandy isthmus at Denpasar.

2.8.  Lombok The north coast of the island of Lombok has cliffs cut in  lava and ash from the Mount Rinjani volcano, and off  the east coast coral reefs are extensive in coastal waters. Beaches are generally of dark grey volcanic sand, often with coralline gravel, and backed by prograded beach ridge plains. Off the northeast coast Gili Petangan, Gili Sulat and Gili Lawang are sand cays on coral reefs, the  sandy beaches containing layers of beach rock (>Fig. 19.8.17), the cays having grassy and shrubby vegetation. On the south coast headlands of sandstone exposed to ocean swell from the southwest are cliffed and fronted by subhorizontal shore platforms produced by weathering down to the water table (>Fig. 19.8.18). There are broad prograded sandy beaches, especially behind coral reefs as at Kuta on the central south coast.

⊡⊡ Fig. 19.8.17 Beach rock on the shore of Gili Petangan, Lombok. (Courtesy Geostudies.)

19. 8

2.9.  Lesser Sunda Islands Many of the features seen on Bali and Lombok are repeated on the Lesser Sunda Islands to the east. High cliffs of ­limestone and volcanic rock fringe the southern coasts of Sumbawa, and Sumba, and active volcanoes have deposited material on the coasts of Flores and Halmahera. Uplift is indicated by emerged coral reefs, attaining 700 m on Sumbawa and over 1,200 m on Timor, where the sequence dated by Chappell and Veeh (1978) indicated uplift rates of up to 0.5 m per thousand years. Many of the smaller islands of eastern Indonesia are either high volcanic islands or uplifted coralline structures of various kinds.

2.10.  Irian Jaya The north coast of Irian Jaya is generally steep, but rivers have built deltas and beach-ridge plains. Evidence of continuing tectonic activity is common: earthquakes cause landslides on coastal slopes, and the Mamberamo Delta has been modified by subsidence (>Fig.  19.8.18). The south coast borders the broad, swampy lowlands traversed by rivers that open into widening muddy estuaries with strong tidal currents; tide ranges here are up to 6 m. The

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Indonesia

⊡⊡ Fig. 19.8.18 Zone of subsidence across the Mamberamo Delta in northern Irian Jaya, showing the Rombebai Lake formed as the land subsided and the spread of mangroves back into the lowered area. (Courtesy Geostudies.)

rivers have not built protruding deltas, but deposition is locally advancing the swampy shorelines.

References Bird ECF, Ongkosongo OSR (1980) Environmental changes on the coasts of Indonesia. United Nations University Press, Tokyo, Japan Chappell J, Veeh HH (1978) Tectonic movements and sea level changes at Timor and Atauro Island. Bull Geol Soc Amer 89:356–368

Hollerwöger F (1964) The progress of the river deltas in Java. Scientific problems of humid tropical zone deltas and their implications, UNESCO, Paris, pp 347–355 Symons GJ (ed) (1888) The eruption of Krakatoa, and subsequent phenomena. Royal Society, London Tjia HD (1975) Holocene eustatic sea levels and glacioeustatic rebound. Z Geomorpholo Suppl.(Bd) 22:57–71 Verstappen H (1973) A geomorphological reconnaissance of Sumatra and adjacent Islands. Wolters-Noordhoff, Groningen

19.9 East Timor

Wong Poh Poh

1.  Introduction East Timor is the eastern half of the island of Timor in the Indonesian archipelago, with an outlier known as Oecusse on the northwest coast. The country became independent in 2002. Much of the country consists of rugged, forested mountains trending west to east, with many deeply incised valleys opening to both north and south coasts. Much of the north coast is steep, with low-lying areas at valley mouths, but there is a more extensive coastal plain in the south. Hard metamorphic rocks dominate the northwest, but there are limestone areas in the eastern part of the country. The climate is tropical monsoonal, with northerly winds in the wet season (November–May) when thunderstorms and heavy rain alternate with drier periods of weaker SE winds. In the dry season (June–October), the prevailing winds are the SE trade winds, which however bring rain to the south coast. Waves generated by these winds are generally small, but nearshore waters can be choppy in the SE monsoon. Mean spring tide range at Dili on the north coast is 1.6 m, increasing to about 3.0 m on the south coast.

2.  The Oecusse Coastline The outlier of the Oecusse country in northwest Timor has a coastline about 65 km long, facing NNW. It is a steep coast with promontories separating valley-mouth bays with small deltaic lowlands. East of Ponta Brente is the mouth of the Besi River valley, then Ponta Lisim and a sandy beach along a bay opposite the re-entrant north of N’itibe. Beyond Ponta Batumera, the shore is stony with segments of sandy beach, and east of Ponta Colan Sina the town of Pante Macassar stands behind a beach-fringed bay. The steep coast continues eastward past Ponta Sacato to the Indonesian border.

3.  The East Timor Coast The coastline of East Timor is about 670 km long. The north coast is generally steep, with bold forested bluffs, as

at Ponta Fatu Su, descending to a bouldery or sandy shore, the sea floor dropping steeply away. The Rio Lois valley opens westward to a small blunt delta at Tailaco, the Lois being one of the few rivers that continues to flow through the dry season. North of this valley is the high Guguleu Range, which has steep slopes descending to the north coast. Ponta Parimbala is a bold headland, and there are sandy beaches at Maubara. At Liquissa, two gravelly rivers open to the shore, and another to the east at Aipelo, on small coastal lowlands prone to river flooding. From Kaitehoe Point to Tibar Point, the shores of the Baia de Tibar have a fringing coral reef, and there is a sheltered anchorage at Nova Alges. Landslides are common on steep headlands such as Tibar Point. The Rio Comoron descends to a delta at Ponta de Motaei, and to the east Dili stands on low-lying land behind a bay that is partly blocked by coral reefs running in to Ponta Fatucama. Offshore, to the north, is Atauro Island, well known for its tectonically uplifted coral reef terraces (Chappell and Veeh 1978) (>Fig. 19.9.1). East of Dili, the fringing coral reef is almost continuous along the steep coast to Manatuto, where the valley of the anastomising Rio Lado do Norte opens to a segment of coastal plain. The coast consists of a series of headlands interrupting bays and areas of mangroves (>Fig. 19.9.2). A well-formed beach lies immediately to the east of the headland with the well-known Jesus statue (>Fig. 19.9.3). Fringing reefs resume to the east, interrupted around the mouth of the Rio Loleia and continuing past Ponta Bandura, where the hinterland rises to the limestone and shale plateau of Salazar. Bacau stands behind a bay with a small delta, and the steep coast, descending to fringing coral reef, runs on past Lautem, where there is a series of rocky limestone headlands, such as Ponto Chater, Ponto Apile, which has a natural arch, and Ponto Hero. The ­hinterland has coral reef terraces indicative of intermittent uplift in Quaternary times (Chappell and Veeh 1978), and rises to the Lautem Plateau. The coral sequences have been used to detect past climatic variations (Hantoro et al. 1997). There are rugged limestone tors on coastal slopes. At the eastern end of Timor, a steep cliff faces across a narrow Jaco Strait to a high reef-fringed island, Ilheu de Jaco. On the southeast coast, the steep craggy limestone

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 19.9.1 High sea level terraces accentuated by ­vegetation and a notch at the lowest terrace, Atauro Island.

⊡⊡ Fig. 19.9.2 Combination of headland, bay, and mangroves, Dili to Manatuto.

slopes of the Monte Paitchau Range descend to a reeffringed coast near Lore and Silvicola, with segments of sandy beach past Ponta Su Lore to Ponta Roro Ai. Here, the coastal plain begins to widen westward past Iliamar and Leça. Fringing reefs occur intermittently to Ponta

Béaco, but are interrupted in the muddy shore sector at the mouth of the Rio Cua valley, where mud volcanoes occur. To the west, the shore is generally muddy, with narrow segments of sandy beach fringing a broad swampy coastal

East Timor

19.9

⊡⊡ Fig. 19.9.3 Cristo Rei East Beach, east of the headland with the Jesus statue.

plain traversed by several rivers descending from the mountain ranges. The Rio Lacla has formed a wide blunt delta with sandy shores. There are short segments of fringing coral reef round headlands such as Ponta Meti Boot, in front of the Besusu Plain. Dry season trade winds generate surf along parts of the shore exposed to the southeast. West of Betano, a sandy beach extends past a series of small deltas, such as that of the Rio Belulic and Rio Laumea. Ponta Suai has a fringing coral reef, behind which a spit has grown eastward, but Ponta Tafara is a sandy deltaic cusp. A long curving bay with a narrow

sandy beach fronting a swampy deltaic plain extends westward to the mouth of the Rio Masin at the Indonesian border.

References Chappell J, Veeh HH (1978) Late quaternary tectonic movements and sea level changes at Timor and Atauro Island. Bull Geol Soc Am 89:356–368 Hantoro WS, Narulita I, Sofjan J (1997) Recent climate variation signals from coral in Timor. Quat Int 37:81–87

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19.10 Papua New Guinea

Eric Bird

1.  Introduction

a­ rchipelagoes to the southeast are added, the total is more than 9,000 km. It is well shown on the 1:100,000 map series. Much of the Papuan coast is low-lying, bordering wide swampy plains, but in New Guinea steep coasts are extensive. There is widespread evidence of Holocene tectonic activity, marked by uplifted coral reefs, especially along the north coast of New Guinea, the south coast of New Britain, and the island chain from Manus to Buka. Earthquakes have disrupted coastal features locally, and

The coastal features of Papua New Guinea (>Fig. 19.10.1) are similar to those of Irian Jaya (Indonesia), with extensive steep sectors backed by uplifted mountain ranges interspersed with deltas and coastal plains (Löffler 1977). The coastline is about 5,000 km long, and when the coasts of New Britain (1,200 km), New Ireland (800  km), Bougainville (600 km), the associated outlying  islands to the north, and the D’Entrecasteaux and Louisiade

⊡⊡ Fig. 19.10.1 Papua New Guinea and the Bismarck Archipelago: location map and distribution of selected coastal features. (Courtesy Geostudies.) 100

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_19.10, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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vulcanicity is active on islands off Madang and near Rabaul, where deposition of lava or ash has prograded some sectors of coastline. Coral reefs are extensive, most coastal promontories and high islands having a fringing reef, while a barrier reef runs off the southeast Papuan coast, looping around Tagula Island and back towards the D’Entrecasteaux Islands, where there are various patch reefs and atolls. Cliffs occur locally on the Papuan coast, but bold coasts are generally steep and forested rather than cliffed. Where Holocene progradation has occurred, it has usually been nourished by sediment brought down by the numerous rivers; the sand fraction has been deposited as successive low beach ridges, and the finer silt and clay have formed associated swamps bearing mangrove vegetation. Beach sediment are mainly black or grey sands and gravels brought down by rivers from volcanic and metamorphic rocks in the hinterland, the proportion of white coralline material increasing towards eroding coral reefs. Parts of the Papuan coast receive sand from fluvial sediment, but on the north coast there are beach gravels delivered by short, steep rivers draining uplifted hinterlands. Of the larger rivers only the Sepik has built a major protruding delta: the Ramu and the Markham valleys end abruptly seaward, while the Fly, Wawoi, Kikori, and other rivers draining into the Gulf of Papua open into estuarine channels between extensive mangrove-fringed swamps, bac­­ ked  by nipa palms and swamp forest. Coastal dunes are rare, but they occur locally on the south coast, notably at Hood Bay. The coastal climate is generally humid tropical, but a relatively dry winter season becomes more pronounced in Papua. Tropical cyclones do not occur. Waves are mainly southeasterly from the Coral and Solomon seas, and northwesterly through the Bismarck Sea. Wave action is low to moderate, especially where reefs and shoals break up distantly derived swell. Mean spring tide ranges are generally less than a metre,  except along the southern Papuan coast, where they increase to 2.8 m at the head of the Gulf of Papua (>Fig. 19.10.1). Tsunamis of up to 15 m have been recorded on the north coast of New Guinea following earthquakes or sudden subsidence of parts of the continental shelf. Evidence of changing sea levels in this region is difficult to interpret because tectonic uplift has been extensive and variable, but some observers have taken raised beach ridges as evidence of a higher Holocene sea level. Human impact has been relatively slight, except where artificial structures have modified shoreline features, in some cases leading to beach and shoreline erosion.

2.  The Coast of Papua East from the Irian Jaya border a broad, lobate lowland with swamp forest and a mangrove fringe extends to Health Bay at the mouth of the Morehead River. Beyond this, narrow sandy beaches fringe swampy lowlands on the coast of Torres Strait, past Walater Point and Maguan Point, to the mouth of Wassi River and the larger Mai River. The Australian border is nor far offshore, and the swampy Kawa Island, Boigu Island, and Saimai Island belong to Australia. Mangroves extend around Buiamuba Point and the mouth of the Oriono River, which reaches the sea in the lee of Bristow Island. The port town of Daru stands on Bristow Island, which belongs to Papua. At Parama Island the coast swings sharply northward on the western side of the Gulf of Papua. The very wide funnel-shaped estuary of the Fly splits between elongated deltaic swampy islands and mangrove shoals, and has tide ranges of up to 2.4 m, accompanied by strong ebb currents. Extensive progradation is indicated by beach ridges stranded within the swamps up to 20 km inland. The swampy deltas are intersected by networks of meandering channels, widening seaward as tidal scour increases. The outer shores of the deltaic islands are lobate, wave action being weak in the shallow nearshore waters. Coastal waters are turbid, with suspended clay and silt brought down by the rivers. A similar swampy coast is seen around the mouths of the Aramia, Turama, and Kikori rivers to the north, but there are small sandy beaches locally, and a spit has formed at Cape Blackwood. Sand brought down during river floods forms elongated intertidal shoals in Deception Bay, and these are gradually reworked and driven onshore by southeasterly wave action. Mangrove-edged swamps dominated by the Sonneratia, Rhizophora, and Bruguiera species occupy the intertidal zone around the head of the Gulf of Papua. Within the mangrove area the surface is made irregular by mounds 2–5 m wide and up to 1.5 m high, built by crabs. Sandy beaches occur intermittently between tidal inlets and river mouths along the shoreline of the Purari Delta (>Fig. 19.10.2). The sandy beaches become more continuous in Orokolo Bay, where the predominant southeasterly waves generate westward longshore drifting. The coastline has prograded with the addition of numerous beach ridges, and to the south the Aumu spit has grown from the sandy mouth of the Vailala River in stages indicated by successive landward recurves (>Fig. 19.10.3). Kerema Bay is flanked by sand spits at Ipisi Point, and beaches continue to the southeast, with mangroves around

19.10

Papua New Guinea

⊡⊡ Fig. 19.10.2 Coastal features of the Purari Delta, where sand supplied by rivers is reworked by wave action and deposited as beach ridges on the seaward margin of a prograding swampy coastline. (Courtesy Geostudies.)

tidal inlets. Progradation of up to 8 km is indicated by beach ridges at Malalaua, where the coastal plain has had a complex history of fluvial sand supply, beachridge building, and subsequent truncation. To the south, beach drifting has influenced the northward growth of the elongated sandy foreland at Iokea. South of the Iokea foreland the hinterland is hilly and cliffs line the coast, interrupted by incised river valleys. The steep coast continues round Cape Possession, behind Hall Sound and the steep-sided Yule Island, and Lagada Island with a cliff at its northern end, Redscar Head. Cliffs and bluffs truncate the hilly ridges that run NNW–SSE to a coast of small bays and headlands, then to a hilly promontory that shelters Port Moresby Harbour (Manala Gadona). Port Moresby stands on the eastern side of this bay, where the coast runs down to Era Point. Coastal waters are less turbid here, and an intermittent coral barrier reef extends southeast along the coast of Papua.

Southeast of Port Moresby the coast is steep and locally cliffed, with occasional river mouths and inlets fringed with mangroves. The larger rivers have carried large quantities of sandy sediment, transported rapidly downstream from steep catchments, to reach the coast as a partly weathered mixture of lithic fragments and various minerals, with only small proportions of quartz. Sectors exposed to swell from the Coral Sea have long, sandy beaches on gently curving alignments, and beach-ridge plains curve out to cuspate forelands, related to wave refraction around reefs. Gabagaba is a village built on stilts in front of mangroves and swamp forest. The barrier reef touches the coast at Hood Point, and to the east is Hood Bay, where sand from the Kemp Welch River has been built into parallel dunes, backed by parabolic dunes up to 35 m high, all well vegetated (Löffler 1977). Hood Lagoon is fringed by swamp forest, and has a narrow outlet between spits.

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h ac be

⊡⊡ Fig. 19.10.3 Sandy depositional features around the mouth of Vailala River, which replenishes shoals during each flood. The sand is being carried northwestward and added to the growing recurved Auma spit. (Courtesy Geostudies.)

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The embayed coast continues southeast to Keppel Point and Paramana Point, and the outlying barrier reef has occasional cays, as at Coutance Island. In the next bay McFarlane Harbour occupies the swamp-fringed Marshall Lagoon at Kupiano, and the hilly coast then runs southeast past a succession of headlands and bays (Kaligala Point with the Brethren Isles, Cheshunt Bay, Inawaimana Point, cuspate Lalaura Point, Kini Kini Bay, and lobate Cape Rodney). Sandbank Point is a cuspate foreland, and mangroves extend behind Cloudy Bay. Dedele Point is cuspate, built out on to a coral reef, and Henderson Bay curves out to another cuspate foreland, Buruma Point. Batumata Point is a sandy foreland with a succession of beach ridges, indicating that there has been progradation on its southeastern shore. In Table Bay a sandy beach fronts swamp forest out to Onibu Point, and Amazon Bay has a beach ridge plain on either side of the Bailebu River. The sandy coast runs out to a cuspate foreland, Kwaipomata Point, built in the lee of an offshore coral reef and cay, Loupumu Island. East of Amazon Bay a steep irregular coast extends past Mapri Bay to Millport Harbour. There are fringing reefs on headlands, as at Gedaisu, as well as the outlying

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barrier reef segments. Orangerie Bay has sandy shores backed by mangrove areas, and Debauo Point protrudes from a mangrove coast on the western side of Mullins Harbour, a landlocked swamp-fringed bay. On the southern side of Mullins Harbour is a high island, Bona Bona, off the western end of a steep hilly peninsula, the southern coast of which is deeply embayed. Between Mullins Harbour and Samarai the coast is exposed to relatively strong wave action through gaps in the barrier reef, and deep bays have not yet been filled with fluvial sediment because their catchments are small and lowlying. Cliffed headlands and bays continue along a hilly coast to Baxter Harbour and behind Gouagurina Bay, where the Owen Stanley Range becomes a mountainous peninsula that ends in a chain of islands off Samarai. Beyond, to the ESE, are the numerous islands and reefs of the Louisiade Archipelago, including several that have been uplifted. The south coast of Misima Island has seven emerged coral terraces up to 460 m, indicating episodic uplift. Milne Bay is a steep-sided embayment heading back to a deltaic plain built by the Gumini River, with a high ridge out to East Cape on its northern side. From East Cape steep coasts run westward behind Goodenough Bay, at the

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head of which the Ruaba River descends to cross a deltaic plain. A hilly peninsula then runs out past Baniara to Cape Vogel, protruding into Ward Hunt Strait. Beyond this strait the outlying high D’Entrecasteaux Islands are accompanied by many reefs and atolls, some with cays. To the north, emerged coral terraces are seen in the Trobriand Islands, whereas Woodlark Island to the east has a tilted coral platform encircling a hilly volcanic core. West from Cape Vogel a hilly coast is fronted by a mangrove zone, then coral reefs on the shores of Col­ lingwood Bay. Rivers are incised into a series of terraces, and there are deltas built by the Kwagira and Rakua rivers. Beyond Benanda Point the coast becomes low-lying and swampy, and the volcanoes of Mount Victory (1,925  m) and Mount Trafalgar (1,644 m) rise steeply on the western side of Collingwood Bay. Mount Trafalgar is dissected by a radial pattern of incised valleys, so that the coastline is intricate where the sea has risen to submerge numerous valley mouths and form deep rias, such as Amuian Bay and MacLaren Harbour, between steep interfluvial headlands prolonged by fringing reefs. One of these runs out on the northern side to Cape Nelson. The Musa River flows down to a large swampy deltaic plain on the southern coast of Dyke Ackland Bay, west of Cape Nelson. The western coast has a narrow coastal plain backed by slopes rising to the Mount Lamington volcano. There was a major eruption here in 1951, when mudflows and torrential runoff delivered masses of sediment to the nearby shores. Beyond Buna the coastal plain widens the Samboga River delta at Cape Sudest, and in Holnicote Bay a sandy beach ridge plain is backed by broad swampy terrain traversed by the Kumusi River. A hilly areas runs out to Cape Ward Hunt, and the coast swings west behind Mambare Bay, where the Mambare River crosses a deltaic plain at Manau.

3.  The New Guinea Coastline The continental shelf narrows past Cape Ward Hunt, and Hercules Bay receives northeasterly waves from the Solomon Sea. The Wama River delta has sandy shores, and Eware Inlet is almost enclosed by a low peninsula, which runs to a barrier spit sheltering a lagoon, Morobe Harbour. The coast becomes hilly and indented below the Bowutu Mountains northwest of Morobe, with Braunschweig Harbour in one of the small bays. The southern shores of Huon Gulf show reef-fringed hilly promontories alternating with low, notched cliffs on raised coral between Morobe and Salamaua, and embayments where streams open through mangrove swamps.

Lasanga Island is steep and high, separated from the mainland by the Royle Channel, and the hilly, indented coast runs on to Salamaua, where Parsee Point is at the end of the peninsula formed by a high island attached to the mainland as a tombolo. The coast continues past Samoa Harbour to Schneider Point, where the coastal hills begin to recede inland. The River Markham has a wide, gravelly watercourse draining a broad valley that opens through a swampy area to a coastline fringed by a beach of sand and gravel. A submarine canyon heads a short distance offshore. To the north, Lae is built on a gravelly river terrace truncated seaward by bluffs. The coast runs eastward along Huon Gulf, dominated by features produced by river deposition, such as alluvial fans and confluent deltaic plains truncated along a coastline where the sea floor plunges to deep water. The Busu River is one of several gravelly streams that end abruptly at the coast (>Figs. 19.10.4 and > 19.10.5). ⊡⊡ Fig. 19.10.4 The mouth of the River Busu, discharging into Huon Gulf. (Courtesy Geostudies.)

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⊡⊡ Fig. 19.10.5 Gravelly shoal at the mouth of the River Busu. (Courtesy Geostudies.)

At Cape Arcana a looped barrier encloses Mundala Lagoon, and at Cape Gerhards the coast is set back to the mouth of the Mongi River. The Gingala Islands are emerged coral reefs offshore. At Cape Cretin the coast swings north to Finschhafen. A spit has grown northward as the Nugidu Peninsula. The coastline curves past Kitumala Point and runs northwest from Fortification Point. On the north coast of the mountainous Huon Peninsula emerged coral reef terraces rise from sea level west of Sialum, on the southern side of Vitiaz Strait. The emerged reefs alternate with truncated alluvial fan deposits, and then the terraces of coral and associated deltaic gravels ascend in level and multiply to more than twenty across a zone of late Quaternary uplift, reaching a maximum height of over 600 m above sea level near Sialum. They are deeply incised by gorges cut by the Tewai, Kowangam, Sapara, and Kwama rivers (>Fig. 19.10.6). Chappell (1974) has calculated rates of uplift of up to 30 cm/century on reefs up to 250,000 years old here. There have been many interpretations of sea level and climate history from these terraces (Yokoyama et al. 2001). The terraced slopes decline past Saidor and Bibi to Astrolabe Bay, where the coast turns sharply northward past the Nuru River delta to Madang. The Madang coast shows emerged coral reefs, cut back on their seaward margins. Emerged reefs diverge northward to Sek Island. At Sarang Cape Croisilles faces across Isumrud Strait to Karkar Island, one of a series of steep conical volcanic islands that runs eastward to New Britain. They include Long Island, which contains a crater lake, Lake Wispom,

about 200 m above sea level, containing a bare rocky island which is an active volcano. Northwest of Sarang a steep coast with coral terraces continues past Cape Gourdon to Awaro, where Manam Island, also volcanic, lies seaward of Stephan Strait. The steep coast ends abruptly, and swampy deltaic plains extend past the truncated Ramu Delta westward to the larger Sepik Delta, where the meandering river reaches the coast at Cape Girgir. The almost straight northern flank of this delta shows evidence of sub­ sidence, where the Murik Lakes are meanders of a former Sepik distributary enlarged by submergence (>Fig. 19.10.7). A sandy fringe is migrating landward into mangrove swamps and nipa forest around the Kaup Lagoon as this coastline recedes. Much of the north coast of New Guinea, west from the Sepik Delta, is steep on the flanks of mountain ranges. The steep Hansemann Coast (>Fig. 19.10.8) west of the Sepik Delta, flanking the Alexander Range, consists of vegetated bluffs frequently scarred by landslides and shows only minor basal cuffing. Offshore, the Schouten Islands are a chain of small volcanoes, some of which are active. The steep coast continues past Richthofen Point and Terebu on Nightingale Bay to Cape Terebu, Foroki Point, and Dove Bay to Cape Moem, the first of a series of emerged coral reef headlands. At Wewak emergence has attached several former coralline islands to the mainland as promontories (tombolos), now with modern fringing reefs (>Fig. 19.10.9). The intervening bays are backed by beach ridges, which curved

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⊡⊡ Fig. 19.10.6 Coral terraces preserved on interfluvial ridges between deep river gorges near Tewai. (Courtesy John Chappell.)

⊡⊡ Fig. 19.10.7 Entrance to the Murik Lakes. (Courtesy Geostudies.)

out to cuspate forelands in the lee of the former islands. The beaches in Boram Bay, Kreer Bay, and Wewak Bay weret formerly supplied with sand by longshore drifting from the west, but they are now eroding, relict beaches, no longer receiving sand from the west. This is because Cape Wom, another promontory formed by emergence and the attachment of a former island to the mainland by a swampy isthmus, has become a natural breakwater, intercepting eastward drifting (Bird 1981).

To the northwest the beach runs out as a lobate barrier enclosing a lagoon on Cape Pus (>Fig. 19.10.10) on the southern side of Muschu Passage. Muschu Island, offshore, is a low-lying emerged coral reef, but Kairiru Island, which lies beyond, is a high island rising to 760 m. The coastline steepens westward past Cape Karawop, with segments of narrow coastal plain, as at Dagua, with beach ridges and swamps. Several rivers descend steeply to

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⊡⊡ Fig. 19.10.8 The steep Hansemann coast. (Courtesy Geostudies.)

⊡⊡ Fig. 19.10.9 Coastal landforms in the Wewak area. (Courtesy Geostudies.)

the coast, delivering sand and gravel to the eastwarddrifting beaches: they include the Ninahau River. Offshore, the Tarawai and Walis Islands are emerged coral reefs.

West of Cape Djeruen the steep coast ends as the Torricelli Mountain front runs inland behind a narrow deltaic coastal plain, crossed by the Dandriuad, Harech, Drinumor, and Nigia rivers. Sand and gravel delivered

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⊡⊡ Fig. 19.10.10 Cape Pus and its lagoon. (Courtesy Geostudies.)

⊡⊡ Fig. 19.10.11 The wide sandy beach at Aitape. (Courtesy Geostudies.)

by  the swift-flowing rivers is carried eastward along the  coast by the prevailing northwest waves from the ­Philippine Sea. In 1935 a major earthquake in the Torricelli Mountains caused landslides and catastrophic flooding, briefly augmenting the yield of sediment to the coast. To the west Aitape (>Fig. 19.10.11) stands in the shelter of a hilly promontory and has a wide beach that has prograded in stages. Towards the rocky headland to the west the sandy beach descends to a boulder-strewn

shore  (>Fig. 19.10.12). Small headlands separate sandy beaches  shaped by northerly swell towards Rohm Point (>Fig. 19.10.13). West of Aitape is a wide, swampy, coastal plain, which includes the Sissano Lagoon, formed when the area subsided 2–4 m during a 1907 earthquake, leaving a narrow, sandy barrier in front of a submerged swamp. The lagoon has been perpetuated by further episodes of earthquake subsidence, and the barrier was overwashed by a tsunami in 1998. There was a major tsunami here in July 1998,

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⊡⊡ Fig. 19.10.12 Bouldery foreshore fronting cliff at Aitape. (Courtesy Geostudies.)

⊡⊡ Fig. 19.10.13 Sandy beaches shaped by ocean swell near Rohm Point, Aitape. (Courtesy Geostudies.)

f­ ollowing an undersea landslide. Evidence of tectonic uplift includes a mangrove swamp deposit raised 50 m above sea level. West of the Biri River delta a coastal range, the Serra Hills, rises and forms a steep coast with cliffs 3 m high at  Prittwitz Point. The outlet from Leitre Lagoon passes through a gorge between high hills at Ana. To the west a narrow coastal plain with beach ridges is backed by swamps. The Nemayer River flows to the sea across a swampy plain with beach ridges, and to the west hilly country comes to the coast at Vanimo, where Cape Con­ cordia is a reef-fringed promontory similar to the attached islands at Wewak. Dakrira Bay to the west is succeeded by another reeffringed promontory out to Vanimo Point, beyond which the coast steepens again along the seaward slopes of the Oenake Range, where emerged coral reefs form terraces up to 100 m above sea level. These descend to low cliffs, up to 2 m high, and an almost continuous fringing reef runs on to the Irian Jaya border.

4.  Bismarck Archipelago The south coast of New Britain is bordered by emerged coral reefs of varying altitude, with terrace stairways near Kandrian similar to those seen on the Huon Peninsula. In  Jaquinot Bay there are up to six terraces, the highest attaining 500 m above sea level. Volcanoes dominate the steep north coast, especially at Cape Gloucester, on the Talasea Peninsula, and around Rabaul, with alluvial fans in

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intervening areas. Coastal deposition of lava and ash, or of volcanic sediment delivered by rivers, has limited coral reef growth here. The southern part of New Ireland has mainly steep coasts, but segments of emerged coral reef occur along much of the north coast and the south coast west of St. George’s Channel, and on some sectors there are multiple terraces up to 400 m above sea level. On the north coast are paired notches along the shore, one at present level, the other indicating Holocene emergence; the elevation of the upper notch increases to the southeast, indicating tilting. New Hanover is almost encircled by emerged reefs, and they also occur on outlying volcanic islands northeast of New Ireland, and on parts of the coast of the Admiralty Islands. Bougainville is largely volcanic with steep coasts, but rivers carrying sand and gravel derived from the volcanoes have supplied material for several segments of beach ridge plain, including Moila Point, a major cuspate ­foreland shaped in the lee of Shortland Island to the south. Fallout of ash may have impeded coral reef development

on the east coast. Hinterland mining operations have augmented the sediment yield from the Jaba River and accelerated coastal deposition (Brown 1974). On the north coast, and especially on Buka Island, a former barrier reef has been raised up to 100 m above sea level, the shore at Cape Putputun consisting of basally notched cliffs behind a modern fringing reef.

References Bird ECF (1981) The beach erosion problem at Wewak, Papua, New Guinea. Singapore J Trop Geogr 2:9–14 Brown MJF (1974) A development consequence: disposal of mining waste on Bougainville, Papua New Guinea. Geoforum 18:17–27 Chappell J (1974) Geology of coral terraces, Huon Peninsula, New Guinea: a study of Quaternary tectonic movements and sea-level changes. Bull Geol Soc Am 85:553–570 Löffler E (1977) Geomorphology of Papua New Guinea. Australian National University, Canberra Yokoyama Y, East TM, Lambeck K (2001) Coupled climate and sea level changes deduced from Huon Peninsula coral terraces of the last ice age. Earth Planet Sci Lett 193:579–587

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20.0 East Asia – Editorial Introduction Much of the coast of > China is hilly, with coastal plains of varying width and many estuaries and rias formed by marine submergence (Eisma 2005). In the south Hainan is an island with a southern upland of granite and a northern lowland of Pleistocene basalt, which extends on to the adjacent Leizhou Peninsula. An embayed coast with many islands, coral reefs and mangroves extends eastward to Hong Kong, and as the coast curves northward there are many rias between irregular promontories. Qiantang River flows into a widening estuary and Hangzhou Bay, while the Yangtze (Changjiang) River opens to a similar estuary to the north. For several centuries the Yellow (Huang he) River flowed into the sea south of the Shandong Peninsula but in 1853 it changed its course to flow northward into the Bay of Bo Hai. Its former delta is eroding, while a new delta has been built into the Bay of Bo Hai, which has a generally low-lying coast. > Taiwan has Miocene volcanic and sedimentary formations along its east and north coasts, where terraced slopes descend to cliffs and beaches exposed to Pacific Ocean swell. There are large waves during occasional typhoons. The north and west coasts are ore-­ sheltered, and fringed by Quaternary sediment and recent deposits delivered by rivers. There are beaches fronted by extensive intertidal flats. Mean spring tide range on the east coast is about a metre, and on the west coast up to 3 m. > North Korea has a low-lying irregular west coast with many islands, exposed to relatively gentle waves on Korea Bay, where the tide range is up to 6 m. Shore ice forms coastal inlets during the winter. There is a smoother embayed east coast with sandy beaches and some cliffed sectors, exposed to north-easterly waves on the Sea of Japan, with a tide range of less than a metre. During winter shore ice forms.

> South Korea also has indented west and south coasts with many rias and islands, exposed to moderate wave action from the Yellow Sea except during occasional typhoons. In winter the sea usually freezes in inlets north of Kunsan. The east coast, exposed to north-easterly waves on the Sea of Japan, has a narrow beach-fringed lowland backed by steep uplands and interrupted by hilly promontories with cliffed headlands. Mean spring tide ranges are large on the west coast (up to 9 m at Inchon) but much smaller on the east coast (1.1 m at Pusan, diminishing northward). The southern limit of shore ice is generally near Samchok. > Japan is mountainous and has many sectors of steep coast, locally cliffed (with shore platforms), and beaches, often backed by dunes. Erosion has led to the making of artificial structures, particularly tetrapods, on many beachfringed sectors. There are coral reefs (some emerged) and mangroves in Okinawa, to the south. Much of the coastline is irregular, with rias and estuaries between rocky promontories. Marine terraces and raised beaches indicate Quaternary uplift, and earthquakes and volcanic eruptions influence the coasts. Wave action is strong on the east coast, with Pacific Ocean swell and occasional typhoons, whereas the west coast is exposed to waves generated by the SW monsoon over the Sea of Japan. Mean spring tide range is between 0.5 and 1.5 m on the east coast, up to 2.7 m in the Inland Sea, between 2.2 and 4.9 m in southern Kyushu, but only 0.1–0.2 m along the Sea of Japan coast.

Reference Eisma D (2005) Asia, Eastern, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 67–71

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

20.1  China    Chen Jiyu . Liu Cangzi . Dai Zhijun . Yu Zhiying    

1.  Introduction With over 5,000 islands the Chinese coastline is 32,000 km long, with a continental coastline of 18,000 km. The geological structure of the coastal area is rather complex, with interlaced distribution of hilly land and low plains. The coastal topography is undulating, and the coastline rather intricate. There are two distinct coastal zones north and south of Hangzhou Bay. The coastal plain to the north has intermittent hilly land, while the coast to the south is dominated by hills with some small estuarine plains (Chen et al. 1980; Chen and Wang 1981; Wang 1980). The Chinese coastline extends across several different climatic zones, varying greatly in natural landscape (>Fig. 20.1.1). South China lies in the tropical zone and the subtropical zone, where the hot and humid climate permits the growth of coral reefs and mangroves. The mangrove coast is widely distributed in South China, extending north to Fujian Province. The coral reef coast is found in the provinces of Hainan, Guangdong, and Guangxi, mainly composed of fringing reefs, while the South China Sea Islands are mainly composed of atoll reefs. Rivers have played a major role in the development of the coastal plains. The Yellow River is famous for its high sediment concentration. Other rivers, such as the Yangtze, also deliver a large amount of silt to the sea, building broad delta plains and muddy coasts around the river mouths, and depositing silt on adjacent continental shelves. The silt from the Yangtze River is transported to the south by longshore currents in Jiangsu Province and Zhejiang Province, and has contributed greatly to the formation of muddy bays in Zhejiang Province and Fujian Province. The muddy coast, with a total length of more than 4,000 km, about a quarter of the total coastline, is mainly distributed along Liaodong Bay, Bohai Bay, north of Jiangsu Province, the estuary of the Yangtze River, bays in Zhejiang Province and Fujian Province, and the estuary of the Pearl River. In addition, sandy or gravelly coasts are found on cliffed coasts and near the mouths of short, steep rivers draining mountainous areas.

Influenced by topography, the tides are rather complex. In most coastal areas they are regular semi-diurnal, but some are irregular semi-diurnal or diurnal. The average tide range varies greatly, from 0.4 m to more than 5 m. Most of the Chinese coast is mesotidal (tide range of 2–4 m), but in Zhejiang Province and areas nearby it is macrotidal (tide range larger than 4 m), and a tidal bore occurs in the estuary of Qiantang River. Guangdong Province and some coastal areas of the Bohai Sea have microtidal coasts (tide range smaller than 2 m). There is a typical monsoon climate on much of the coast of China, northerly winds prevailing in winter and southerly winds in summer. Wave directions are related to these winds, and show seasonal variations, playing an important role in the movement of silt alongshore. Strong waves are mainly caused by typhoons, and are also an important cause of coastal erosion in China.

2.  The Coastline of China Features of the Chinese coastline are described in a sequence from south to north. From the Vietnamese border the Southern Guanxi hilly coast, between the mouths of Dafeng and Beilun Rivers, is mainly composed of Palaeozoic and Mesozoic sedimentary rocks. The coastline is hilly and much indented, with estuaries penetrating into the hinterland. Mangrove forests grow around the bays. The coast turns southward along the Leizhou Peninsula, which consists of platforms at altitudes of 20–80 m above sea level. The northern coast is composed of Neogene and early Quaternary clastic rocks, while the southern coast is formed of Pleistocene basaltic lave flows. The western coast is relatively straight, but the rest of the coast is irregular, indented with drowned valleys and bounded by cliffs. The beach is mainly composed of sand and pebbles, and mangroves grow on the tidal flats. Off the southern end of the peninsula are emerged coral reefs 1.5–2.0 m above low tide level, dated 7,120±165 years bp. To the south is Hainan Island. Its northern coast is dominated by a succession of lave flows, with extensive

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 20.1.1 Coastal geology and landforms of China.

pebble and sand beaches and many fringing coral reefs. A delta has been built by the Nandu River. Mangrove forests of various species cover the bay area, the taller ones being as high as 10 m. The southern coast of the island is mountainous and mainly granitic. The embayments are often enclosed by spits to form lagoons. Beaches and spits are capped by sand dunes parallel to the coastline. There are coral reefs offshore, some emerged.

The South China Sea is dotted with numerous small islands in addition to Hainan. They can be classified into four groups: the Dongsha Islands, the Xisha Islands, the Zhongsha Islands, and the Nansha Islands. All are of coralline origin. Since the Miocene atolls have gradually evolved from submerging fringing and barrier reefs, these reefs include skeletal coral and shells. The coral reefs are up to 1,000 m thick, as a result of sea floor subsidence.

China

20.1

⊡⊡ Fig. 20.1.2 Sandy coast on Hainan Island.

In the Holocene some reefs have been raised; the 4,000 year old limestones of Rock Island, for example, are now 15 m above sea level. East of the Leizhou Peninsula is a coastal plain. Rivers flowing into the South China Sea are usually fed by the confluence of numerous torrential streams originating in mountainous regions. Though estuarine accumulation has been taking place since the middle Holocene, the alluvium has grown only slowly because of decreasing sediment supply. The Zhujiang estuary and the Hanjiang River ­estuary (the largest in southern China) still maintain their original triangular shape where a valley mouth has been invaded by the Holocene marine transgression. The Shaiang River delta is formed by the confluence of three subdeltas, the Xijiang, the Beijiang, and the Dongjiang deltas, with the sediment supply from five rivers including the Xijiang, the Beijing, the Dongjiang, the Liuxi, and the Tanjiang. The deltaic plain, dotted with isolated hills, is drained by a network of interwoven channels emptying to the sea through eight outlets. Progradation, related to the location of these outlets, varies from 10 m/year to 110 m/ year in response to changes in runoff and sediment supply. Broad mudflats have formed on sectors where the accretion process is relatively active. The Hanjiang delta is fan-shaped. The river flows into the sea through three distributaries. The delta is growing seaward at the rate of 10–30 m/year, and numerous sand ridges have been formed by marine reworking of fluvial sediment.

Hong Kong has a highly indented coastline with many peninsulas and islands (So 1985). It consists of granitic and volcanic rocks, exposed on sectors exposed to strong wave action, particularly during typhoons. There had been extensive land reclamation on the north shore of Hong Kong Island and the adjacent mainland in strongly urbanised areas. There are many small bay-head beaches, and the more sheltered inlets contain mudflats and mangrove swamps. The Southern Fujiang and Eastern Guangdong hilly coast is composed mainly of granite. Its weathering products provide abundant sediment, which is reworked by wave and current action. This coast is rocky and sandy, with mudflats and mangrove swamps in sheltered bays and estuaries, such as the Mingjiang and Jiulong Jiang estuaries. The tidal zone has three types of beaches: rock beach, sand beach, and tidal flat (>Fig. 20.1.2). The Eastern Zhejiang and Northern Fujian hilly coast can be categorised as cliffed, sandy, silty, and mangrove coasts. Cliffs, forming more than half the coastline, are cut in Mesozoic volcanic rocks and granite, and are found in southern and northern Zhejiang and on island coasts exposed to high wave energy. Sandy coasts are limited to about 4% of this sector. Finer sediment is found in embayed shores, with intersecting networks of channels, believed to have been formed by strong tidal currents. Due to the relatively weak wave action fine-grained sediment from the Yangtze estuary is carried into the bays by the flood tide, leading to the development of wide muddy marshes and

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⊡⊡ Fig. 20.1.3 Sea wall, Hangzhou Bay.

⊡⊡ Fig. 20.1.4 Changjiang (Yangtze) estuary.

China

intertidal flats. Mangroves grow in sheltered muddy environments on the embayed shore in northern Fujiang. The coastal sector from the southern shore of Hang­ zhou Bay to the Beilun estuary in Guangxi province is hilly, controlled by faults trending NE and MW. The region is divided into hilly coast, river deltas, coastal plain, and intertidal zone. The irregular coastline has many natural harbours, peninsulas, and islands (Dai et al. 2004). Small coastal plains occur beside estuaries and behind bays. Hangzhou Bay is macrotidal, with an average tide range exceeding 5 m. In the upper reaches of the Jinshan spit the tide range is over 4 m, and up to 5.62 m in Ganpu. The bay retains its funnel shape because the silt supply from the Qiantang River has been insufficient to fill it. Sediment on the floor of Hangzhou Bay comes mainly from seaward sources, and amounted to 38 billion m3 from 1959 to 2003. The average annual fluvial sediment yield is only 5.4 million tons. The rapid narrowing and abrupt shallowing of the bay headward have deformed incoming tides to generate tidal bores, such as the famous Qiantang bore west of Ganpu. The tidal bore in Yanguan may reach as high as 3.7 m, the maximum current velocity of the flood tide being 16–20 kts. Apart from a few isolated coastal hills Hangzhou Bay is backed by a coastal plain. The intertidal flats are silty. Over a long period the strong tidal currents of Hangzhou Bay, combined with the lateral movement of the deep channel, have resulted in erosion and accretion. While the coastline on one side was rapidly silted up and hence prograded, the other was strongly

⊡⊡ Fig. 20.1.5 Eroding tidal mudflats, central Jiangsu.

20.1

scoured and hence cut back. At present the northern bank is subject to slight erosion, while the southern bank shows a tendency toward accretion. Sea walls have been built to protect eroding sectors (>Fig. 20.1.3). The Yangtze River, the longest river in China, delivers a large volume of runoff and sediment to the sea. Its annual sediment yield is over 4.7×108 tons. It forms a large deltaic area with a submarine delta offshore. The Yangtze estuary was a funnel-shaped bay 2,000 years ago, but rapid deposition has formed a delta, with river channels first being divided by Changming Island into the North and South Branches, then the South Branch being split into the North and South Channels by the Changxing and Hengsha islands, and, finally, the South Channel branching into the North and South Passages on either side of Jiuduansha Shoal (>Fig. 20.1.4). The outlines continue to change, and the North Channel has been deepened by dredging (Chen et al. 1959, 1979). The Jiangsu Plain coast from Haizhou Bay to the Yangtze Estuary is a muddy lowland, except in the Lianyungang region, where local hills have developed due to fault block uplifting. The four sand ridges on the coastal plain are indicative of the south-north shifting of the ancient Yellow River. Between the years 1128 and 1855 the Yellow River flowed out through the lower reach of the Huai River into the sea. A huge amount of silt was supplied to build a delta centred on the now abandoned Yellow River estuary. In 1855 the river migrated northward again, and the coastal section north of the Sheyang River then

1193

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20.1

China

retreated because of the abrupt loss of ­river-borne sediment supply, and exposure to wave attack. The recession rate reached 150–200 m/year, and the tidal flat sediment became coarser as mud scarps and shell ridges were formed at the seaward edge of the marshes (>Fig. 20.1.5). The coastal section south of the Sheyang River continued to prograde because offshore shoals (both emerged and submerged) protected the coastline. In addition, silt from the Yangtze River was deposited here. A series of offshore sand ridges radiate outward in finger-like patterns produced by the interaction of tidal currents from the south and the north. Thus the newly-deposited ridges arose from the reworking of surface sediment of the ­subaqueous delta by tidal currents. In central Jiangsu, Spartina alterniflora

resulted in accretion action after it was planted on the tidal mudflat (>Fig. 20.1.6). Spartina alterniflora has been replaced by Spartina anglica in the area since 1980. The Shandong Peninsula extends from Hutou-ya Town to the Lanshantou Hills, and terraces are composed mainly of Pre-Cambrian metamorphic rock and granite intrusions. The south coast of the peninsula is indented, with capes and bays including Laoshan Bay and Jiaozhou Bay. The north coast is simpler, shaped by waves from the Yellow Sea, with sandy beaches and tombolos, as at Zhifu Island. There has been some beach erosion (>Fig. 20.1.7). Laizhou Bay, to the west, has a 4–6 km wide muddy intertidal zone, with silt mainly from the Yellow River. The ⊡⊡ Fig. 20.1.6 Spartina marsh, north shore of Hangzhou Bay.

⊡⊡ Fig. 20.1.7 Coastal erosion, Shandong Peninsula.

China

20.1

⊡⊡ Fig. 20.1.8 Erosion of abandoned Huanghe delta, Shandong.

⊡⊡ Fig. 20.1.9 Rocky coast on Liandong Island.

Yellow River delta is a 200-km-wide sector, protruding into the shallow Bohai Sea. The river delivers large quantities of sandy sediment, on average 8.7 × 108 tons annually, 40% of which is deposited in the deltaic area. The sand spit at the river mouth is growing seaward at a rate of 2–3 km/year. After 1855, when it was diverted back to the Bohai Sea (see above) it has formed a large complex delta, composed of 10 sub-deltas related to changes in its course. The present sand spit at the estuary formed after the river was artificially diverted in 1976 through the Qingshui channel, and has extended 28 km further into the sea. Muddy intertidal shores supplied with silt from the river are generally 3–5 km wide, and may reach a maximum width of over 10 km (Pang and Si 1979) (>Figs. 20.1.8, > 20.1.9 and >20.1.10). Northwest of the Yellow River delta the coast consists of a swampy beach-ridge plain, fronted by a muddy intertidal zone generally 3–5 km wide, but locally up to 10 km. Tianjin New Port is a large sea port on the muddy intertidal shore. Further north is the Luanhe deltaic coast. The Luanhe River has supplied fine sand and silt to the shore, forming a delta that is exposed to waves up to 4.8 m high, which have built a barrier island and lagoon system at the front of the delta. Due to frequent changes in course of the Luanhe River successive sub-deltas have been constructed. The modern delta has grown since the present course of the Luanhe River was established in 1915. The Liaoxi hilly coast between the Luanhe River, the Xiaoling River, and the Luan River is on Pre-­Cambrian

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China

⊡⊡ Fig. 20.1.10 Cliffy coast on Liandong Island.

metamorphic rocks with granite inselbergs. The straight coastline is fringed by a beach supplied with sand from rivers draining the Yanshan Range. There are up to six ­parallel beach ridges, and a coastal dune 20 km long, 1–3 km wide, and 30 m high has developed on the backshore. Liaodong Bay, in the north of Bohai Sea, has a coastal plain formed by deposition of sand and silt from the Shuangtaizi, Liaohe, Daling, and Xiaoling rivers. Coastal wetlands are fronted by a muddy intertidal zone up to 2 km wide, with a very gentle seaward gradient. East of Liaodong Bay the Liaodong Peninsula is a hilly upland of Pre-Cambrian metamorphic and lower Palaeozoic sedimentary rock, bordered by an intricate embayed coast. The bays have beaches fronted by muddy shores, supplied with silt from the Daze and Yalu rivers, and the intervening headlands are cliffed. Karst topography has developed on limestone outcrops in Liangshui Bay. Sandy beaches are extensive on the more exposed

south coast, facing Korea Bay. The Yalu River estuary marks the border with North Korea.

References Chen JY, Wang BC (1981) The islands of China. In the physical geography of China. Science Press, Beijing, pp 349–365 Chen JY, Wang BC, Liu CZ (1980) The coastal geomorphology of China. In the physical geography of China. Science Press, Beijing, pp 313–347 Chen JY, Yu ZY, Yun CX (1959) Development of the geomorphology of the Changjiang delta. Acta Geogr 24:201–222 Chen JY, Yun CX, Xu Haigen, Dong YF (1979) Evolution of the Changjiang estuary during the past 2,000 years. Acta Oceanol Sin 1:103–111 Dai ZJ, Li CC, Zhang QL (2004) Fractal analysis on shoreline patterns for the crenulate-bay beaches, Southern China. Estuar Coast Shelf Sci 61:65–71 Pang JZ, Si SH (1979) The estuary changes of Hunghe River. I. Changes in modern time. Oceanol Limnol Sin 10:136–141 So CL (1985) Hong Kong. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 823–828 Wang Y (1980) The coast of China. Geosci Canada 7:109–113

20.2 Taiwan

John R. C. Hsu

1.  Introduction The island of Taiwan is separated from the Chinese mainland by the Taiwan Strait, which is about 150 km wide. It has an area of about 36,000 sq km and a coastline about 1,520 km in length (including the Penghu Archipelago to the west). The island is situated on the convergence of the Eurasian and Philippine plates, and the rising of its Eastern Coastal Range is a product of plate tectonics. The island is also part of an island arc at the meeting point of the Ryukyu and Luzon arcs, which bend convexly toward the Asian continent. The eastern half of the island consists of a metamorphic complex of upper Palaeozoic age, bordered by the Eastern Coastal Range of Miocene volcanic and younger sedimentary rock. The western foothills of the Central Range are composed of Tertiary marine sedimentary formations, and the Western Piedmont and the Western Coastal Plain occupy the western half of the island (Ho 1975). The present outlines of the coast are related to these geological formations and to the marine processes operating in the coastal area, as well as to many man-made structures in more recent times. The sea floor in the Taiwan Strait is mostly at depths of 100–150 m, but the sea is deeper towards the Pacific Ocean, declining abruptly to more than 3,000 m a relatively short distance offshore. Taiwan has a subtropical monsoon climate. The winter monsoon from the northeast begins in October and ends in the following April, and the summer monsoon from the southwest lasts from July to September, with moderate rain beginning in May. Wind velocity during the winter monsoons may reach 15 m/s, but much stronger winds blow during typhoons (tropical cyclones) with heavy rainfall, which occur several times during the summer and early autumn, and generate high storm waves that attack the coast, usually lasting for about 24 h (Hsu 1985). Mean spring tide range is up to 4 m on the central west coast, whereas both ends of the island have a range of only about 1 m or less. The central west coast consequently has a broad intertidal zone, ranging from 1–5 km, parts of

which were reclaimed for agriculture in 1960–1980 and for industry after the 1980s. Large quantities of silt and sand and some gravel, derived from the Central Range, were brought down to the coast by rivers during typhoon event. The finer sediment was deposited to form broad tidal mudflats, while the sandy material formed beaches and chains of coastal barriers in the southwestern part of the island (Shih 1980). Longshore drift in response to the persistent waves generated by the northeastern winter monsoon carries sediment southward along the western coast from the north of the island to north of Tainan; however, the drift is predominantly northward on the coast south of Tainan, due to the weakening of the winter monsoon waves arising from the abrupt change in coastline orientation from NE to SE. The persistent longshore drift on the eastern coast is southward, generated by the winter monsoon. The pattern of longshore drift is indicated by accretion of sediment on the updrift side of harbour breakwaters and erosion on the downdrift side.

2.  The Taiwan Coastline The coastline of Taiwan shows contrasts that result from wave action on the coast under various geological controls (Hsu 1962). The west coast between Taisi and Tainan was advancing as the result of deposition and emergence. Much of the sediment was brought down by rivers draining from the central mountains, providing silt and sand for the formation of tidal flats, marshes, and coastal barriers (Shih 1980). Uplift of this area since the Pleistocene contributed to the progradation of the coastline. By contrast, the SW coast between Fengkang and Kaohsung retreated as the result of regional subsidence and marine erosion. Valley mouths had been submerged to form a ria coast, and ridges had been dissected into promontories and islands, some of which were tied by tombolos, as near Suao. On the south coast of Taiwan post-Pleistocene tectonic uplift formed emerged coral reefs, well developed in the Oluanpi area, where the highest is over 200 m above

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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20.2

Taiwan

sea level. Living corals are abundant in fringing reefs along the coast near Oluanpi. Nevertheless, there has been considerable erosion, and much of the coastline is retreating. The coast near Hoping and Tawu, in eastern Taiwan, is straight and simple in outline with continuous sea cliffs, and is thought to have originated as the result of faulting. The eastern flank of the Coastal Range has been uplifted at different rates in different places. As it is exposed to strong wave action from the Pacific Ocean, the emerging coast has been subject to erosion, which has at least partly offset the gain of land by upheaval. The northeast coast near Yilan has been retreating as the result of tectonic subsidence and marine erosion. It has been necessary to protect the farmland of the Yilan plain by stabilising the coastal dune fringe, and by building short groynes to prevent invasion by the sea. The northern coast of Taiwan has also been subject to post-Pleistocene tectonic uplift, but marine erosion has cliffed the emerged fringe. In the Yehliu Scenic Area near Keelung City is an elongated promontory of Miocene sandstone with hoodoo rocks, remnants of concretionary sandstone, exposed by marine erosion (>Fig. 20.2.1). They include mushroom rocks, as at Queens Head, upstand­ing calcareous concretions, as at Ginger Rocks, and small flat-topped conical protrusions, as at Candle Rocks (Hsu 1964).

In recent years damming of rivers for water supply and flood control have intercepted sediment supply from most of the rivers on the island. Excavation of sand and gravel from rivers for construction and land subsidence due to extraction of groundwater for aquaculture have also intensified the problem. Consequently, many parts of the coast on the western coastal plains previously advancing as the result of deposition and emergence, are now eroding. For example, Waisanding, the largest barrier island on the SW coast of Taiwan, has been migrating landward and diminishing in area over the last 20 years. In addition, breakwaters alongside the four major ports and numerous fishing harbours have intercepted longshore drifting sand, resulting in severe beach erosion downdrift. The response has been to build extensive shore protection structures (>Figs. 20.2.2 and >20.2.3). The Water Re­­sources Agency has been responsible for the coastal protection in Taiwan since the 1960s, and hard concrete sea walls and other structures have been installed. By the end of 2007 there were 536 km of sea walls, 59 km of wave dissipation works, 206 groynes, 225 detached breakwaters, and 28 submerged offshore breakwaters. As a result, 51% of the coastline has now been replaced by man-made concrete structures. Since 2005 the government has used artificial beach nourishment, as at Sizihwan, Kaohsiung.

⊡⊡ Fig. 20.2.1 Concretionary sandstone structures (hoodoos) on the shore west of Keelung northeastern Taiwan.

Taiwan

20.2

⊡⊡ Fig. 20.2.2 Pre-cast concrete blocks fronting a sea wall at Kerzihliao in Kaohsiung county.

⊡⊡ Fig. 20.2.3 Detached breakwaters at Kerzihliao in Kaohsiung county have caused leeside accretion and tombolo formation. (Courtesy Taiwan CECI Engineering Consulting Inc., Taipei.)

References Ho CS (1975) An introduction to the geology of Taiwan. Ministry of Economic Affairs, Republic of China, Taipei Hsu TL (1962) A study on the coastal geomorphology of Taiwan. Proc Geol Soc China 5:29–45

Hsu TL (1964) Hoodoo rocks at the Yehliu area, northern coast of Taiwan. Bull Geol Surv Taiwan 15:37–43 Hsu TL (1985) Taiwan. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 829–831 Shih TT (1980) The evolution of coastlines and the development of tidal flats in western Taiwan. Geogr Res 6:1–36 (In Chinese)

1199

20.3  North Korea

D. Eisma

1.  Introduction North Korea has a west coast that extends from the Chinese border at the mouth of the Yalu River south to the estuarine inlet at Haeju-man, and an east coast that extends from just south of Kosong to the Russian border at Kangui. The west coast is indented and hilly, whereas the east coast has a narrow coastal plain fronting the Taebaek and Hamgyong-Sanmaek mountain ranges. The landforms of North Korea have been shaped on Pre-Cambrian crystalline rocks (granite, gneiss and schists) and Late Cretaceous and Tertiary granite intruded into rocks folded and faulted along N–S and NE–SW alignments in Jurassic and Cretaceous times. Miocene and later earth movements caused subsidence on the west coast while raising the east coast. A major N–S fault zone lies in the Sea of Japan off the east coast, forming a submarine escarpment. Because of the tilting, the principal mountain ranges lie near the east coast, so that only short rivers flow to the Sea of Japan. The coast is relatively steep, with a narrow continental shelf (generally not more than 10 km wide) and a rapid descent to the ocean floor of the Sea of Japan at more than 2,000 m depth. The west coast is much flatter and gradually merges into the Yellow Sea, which is a shelf sea less than 100 m water deep. The tilting is also reflected in the distribution of coastal terraces. A high terrace, mostly erosional, with local marine gravels and a cover of red soil and weathered crust, stands at 50–100 m along the east coast, and 20–30 m along the west coast. A middle terrace, with extensive gravel beds but no red soil, stands at 30–80 m along the east coast and at 10–20 m along the west coast. Along the east coast there is also a third terrace at 10–20 m which is absent from the west coast (Oh 1978), unless a locally developed rocky platform, 1 m above, has highest spring tide correlated with it (Guilcher 1976). The climate is temperate, but with cold winters when winds blow from the northwest. Pyongyang has mean monthly temperatures of −7.8°C in January and 23.9°C in July, with an average annual rainfall of 916 mm. Coastal waters freeze for up to 4 months in winter on parts of the west coast, and for a shorter period on the milder east

coast. Tide ranges are small (Fig. 20.3.1). There has been extensive draining and reclaiming of coastal lowlands. Subaerial weathering has resulted in rather steep conical hills and small isolated mountains, separated by broad alluvial plains. Chodo Island is hilly, and the high narrow ridge of the Changam-dong peninsula has one of the few substantial beaches (Samni-dong) on its northern shore, backed by low dunes in an arcuate bay facing north-west. Marine erosion has been weak since the Late Quaternary submergence, but some of the promontories and outlying islands more exposed to wave attack show cliffs and shore platforms. In the West Korean Bay shallow sand ridges up to 80 km long, about 2.5 km wide, and with an amplitude of up to 30 m, have been shaped by the strong tidal currents (Off 1963).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

20.3

North Korea

On the shores of promontories and islands there are many small beaches of variable composition, ranging from fine sand to coarse gravel, and small spits of various shapes and sizes. In bays and between islands, where there is

124°

­ rotection against strong wave action, tidal flats have p grown upward and seaward. Off river mouths they protrude as intertidal deltas. The tidal flats are underlain by Holocene sediment of varying thickness (up to 25 m) and

125°

124°30'

⊡⊡ Fig. 20.3.1 Coastal, nearshore and submarine morphology of the West Korean Bay. (Courtesy Geostudies.)

125°30'

KOREA

CHINA

Yalu river mouth

39°30'

So

su

do

P'yôngyang

39°

To

39°

ng

su

do

39°30'

N

Ta

ed

on

gG

an

o

g

38°30'

0 0

5 10

20 mile

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oS

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sandy beaches salt marches salt pans

Changsan Got

Taedong Man

tidal flats 6 fathom depth line

38°

38°

10 fathom depth line

124°

124°30'

125°

125°30'

⊡⊡ Fig. 20.3.2 Geology and sea floor features, west coast of Korea. (Courtesy Geostudies.)

North Korea

  20.3 1203

20.3

North Korea 127°30'

128°

128°30'

EAST KOREAN BAY Predominant coastal landforms Ya

cliffs, rocky coast sandy shorelines

Yonggo Dan

⊡⊡Fig. 20.3.3 Coastal landforms, East Korean Bay. (Courtesy Geostudies.)

hw

ng a

Hongwon

M

Toegu Dan

an

Sinp'o Hang

40°

Songdo Gap

40°

Mayang Do

Hamhung Ansong Gap Hungnam Oeyangdo Dan Hwa DO Hamhung Man

N

Paegan Dan

TONGJOSON MAN

39°30'

39°30'

Sô Man Songjon Man

Taegang Got Wonch 'u Gak Yônghúng Man Tae Do Sin Do Yo Do Kalma Gak So Do Amyong Kkut

Cho

n

Wonsan

dae

Chasan-ni

Nam

1204

39°

39° Kojo

Ch'ongsok Tan

Sol Som

Chaegdae Kkut 0

38°40'

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10 10

20 20

30

30 miles 40

Suwon Dan

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127°30'

rest in turn on an undulating bedrock surface. There are also emerged Holocene tidal flats, which have been largely reclaimed for agriculture since 1900. The tidal flats are sandy where the coast has been eroded or where tidal ­currents or waves are strong. The upper parts of tidal flats are often covered by periglacial debris (angular pebbles in

128°

38°40'

128°30'

a fine matrix), which has moved downward by solifluction from adjacent coastal slopes. To the south the Hwanghae-Namdo peninsula consists of Ordovician limestone, and has an irregular coastline, with tidal flats in rias between hilly promontories (>Fig. 20.3.2). Submerged valleys can be traced down to

North Korea

42°

42°

cliffs, rocky coast sandy shorelines Granit/gneiss/schists (Precambrium)

Kal Tan

Basalt/rhyolite/trachyt (Pliocene) Basalt (Pleistocene) Ch'ongjin

Komalsan Dan

an H g

Guilcher A (1976) The ria coasts of Korea and their morphological evolution (in French). Ann Geogr 85(472):641–671 Kwon HJ (1974) The intertidal flat of the west coast of Korea and the origin of its sediment (in Korean with English summary). Geography 10:1–12

130°

in gj

References

129°30'

N.E. KOREA coastal geology

on h' C

Kyongsong Man

Canyon

41°

41°30'

'on

Orang Dan

an

Ch

Puam Man Kw ae D Tajin an

gg on

ang

nH

aeji

Od

My

Man Taeryanghwa Man So ng N Da n n Hwangjin Ma Poksuk Tan n k Ta ndo Cho P'ohang Man

41° n

North from Kosong the coast faces northeast and a narrow coastal plain is backed by the rising slopes of the Taebaek Mountains. To the north is an area of subsidence  around the East Korean Bay (the Bay of Wonsan) (>Fig.  20.3.3). Two small sandspits enclosing coastal lagoons form Cape Amyong. Arcuate sandy beaches back the embayment to the north, shaped by swell refracted past the Taegang Gat peninsula, and there are salt marshes on the sheltered northern shores of the bay. A 16-km long sandy spit connects a former island of Pre-Cambrian gneiss with the mainland, the whole forming the N–S Hodo peninsula. Other small islands have been connected to the mainland in a similar way, as on the Galma peninsula, or through the seaward extension of river deltas. The northeast coast has many small and several large bays formed by Late Quaternary marine transgression over a hilly landscape (>Fig.  20.3.4). The rivers in this area, coming from almost barren uplands, carry large amounts of sand and gravel. Where they flow out into a bay, a broad alluvial plain has been formed. The Myonggang River has brought down sufficient sediment to build a lobate delta in a bay that faces north-east. Because of the swell and the coastal current, both coming from the northeast, spits and bars have been built up, but where the coast is rocky and around islands, cliffs have been formed. This part of the coast is considered to have been tectonically relatively stable (Lautensach 1945). To the north the Tumen-Jiang River has built a delta that is dotted with hills and lakes, with the Russian border following the river.

129°10'

Ch'o

3.  The East Coast

⊡⊡Fig. 20.3.4 Coastal geology of north-east Korea. (Courtesy Geostudies.)

41°

Unmandae Dan

Na md ae

depths of 30–50 m at distances of 50–70 km offshore: Below about 40 m they are filled and covered by Holocene mud.

  20.3

Mokchin Dan

Musu Dan Yang Do Kanghui Do 0 5 Yujin Dan

129°10'

0

129°30'

10

10

15 20

20 miles 30 km

130°

Lautensach H (1945) Korea. Koehler Uerlag, Leipzig Off T (1963) Rhythmic linear sand bodies caused by tidal currents. Bull Am Assoc Petrol Geol 47(2):324–341 Oh GW (1978) The characteristics and the development of the shoreline of the Korean peninsula (in Korean with English summary). Geography 18:22–32

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20.4 South Korea

D. Eisma

1.  Introduction The southern part of the Korean Peninsula is mountainous and hilly, with a coastal plain in the southwest. The northeast coast follows a NE–SW trend of faulting and folding that dates from the Jurassic and Cretaceous, and the rectilinear shape of the peninsula reflects a N–S faulting trend dating from the same period. Since Miocene times, the land has been tilted up in the east and northeast and tilted down in the west and southwest. The east coast of the peninsula is therefore a coast of emergence, and the west and the southwest are coasts of submergence, with former inselbergs as islands and promontories. A major N–S fault zone lies in the Sea of Japan off the east coast, along a submarine escarpment. Because of the tilting, the principal mountain range, the Taebaek Mountains, is near the east coast. Only small rivers flow toward the Sea of Japan, and the coast is relatively steep, with a narrow continental shelf (generally not more than 10 km wide) and a rapid descent to the ocean floor of the Sea of Japan at a depth of more than 2,000 m. The west coast is much lower and flatter and gradually merges into the Yellow Sea, which is a shelf sea less than 100 m deep. The south coast, like much of the west coast, is fringed by many islands, whereas the east coast is almost devoid of them. Tilting is also reflected in the levels of coastal terraces in South Korea. Along the east coast an erosional high terrace, with local marine gravel and a cover of red soil and weathered crust, stands at 60–100 m, but it descends to 20–30 m along the west coast. A middle terrace, with extensive gravels and without red soil, stands at 30–80 m along the east coast and at 10–20 m along the west coast. Along the east coast there is a third terrace at 10–20 m which is absent from the west coast, unless a locally developed rock platform 1 m above the highest spring tide should be correlated with it (Guilcher 1976). Pre-Cambrian crystalline rocks (granite, gneiss and schists) and Late Cretaceous and Tertiary granites form most of the South Korean coast, with Cretaceous freshwater deposits in the southeast and the south. Local Palaeozoic and Mesozoic marine deposits and basalts are also ­present.

The island of Cheju-Do, south of the peninsula, and the island of Ullung-Do, in the Sea of Japan, are volcanic, with Cheju-Do having been active in historical times. The climate is temperate but with cold winters when winds blow from the northwest. Pusan, in the south east, has mean monthly temperatures of 2.2°C in January and 24.4°C in July, with an average annual rainfall of 1,407 mm. In contrast to North Korea, coastal waters freeze only briefly and locally in winter, on sheltered parts of the coast. Mean spring tide ranges are generally small along the east coast (1.1 m at Pusan) but much larger on the west coast, where Inchon is macrotidal with a maximum range of 8–9 m during spring tides and more than 3 m at neap tides. There is a large diurnal inequality of 1.41 m, a monthly inequality of about 1 m and a yearly inequality of 35 cm, which is related to variations in water level in the Yellow Sea. There is probably also an 18-year cycle with a range of about 45 cm. Tidal currents are correspondingly weak along the east coast and strong in estuaries, inlets and between islands on the south and the west coasts. Winds are mostly northerly during the winter and variable during the rest of the year. Gales are most frequent in winter, coming from the northwest. On the coast, local winds prevail, including those from nearby mountains and valleys, and in many places, offshore breezes by night alternate with onshore breezes by day. Typhoons occur about once a year during summer, primarily along the southern part of the peninsula, coming from WSW. The east coast receives a NNE swell from the Sea of Japan, but on the south and the west coasts, waves are locally generated, from the Yellow Sea, and their effects are much reduced by the sheltering islands and the larger tide range.

2.  South Korean Coastline From the North Korean border behind Haeju-man, the west coast is indented with numerous rias, between irregular headlands, and many islands. The Han-gang River flows north to an estuary west of Seoul and to the south is the port of Inchon. The large tide range results in the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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20.4  

South Korea

exposure of wide tidal flats at low tide, with some islands linked by intertidal bars. There has been extensive reclamation of salt marshes and tidal flats, and sectors of the coastline have sea walls. The west and the south coasts show evidence of submergence by the Late Quaternary marine transgression and long-continued land subsidence. On the basis of radio­ ­carbon dating of peat, subsidence during the Holocene has been of the order of 0.4 mm/year (Park 1969). There are intricate steep-sided peninsulas and intervening bays, with rivers such as the Geum-gang opening through estuaries to bay heads. Guilcher (1976) described marine beaches at about 1 m above sea level along the south and west coasts as Late Pleistocene (Eemian) and concluded that the coast had remained virtually stable during the Late Quaternary, but these emerged beaches, together with rock platforms, old tidal flats, former cliffs with basal notches at 3 m and old spits at 0.5–3 m above spring tide level, have been taken as indications of a Holocene high sea level at about 1–2 m. Submergence of SW-NE trending hills and low mountain ranges, following one of the main folding and fault directions, has resulted in a typical ria coast, with a series of low peninsulas, large islands and groups of smaller islands farther offshore. Subaerial weathering has resulted in rather steep conical hills and small isolated mountains, separated by broad alluvial plains. Marine erosion has been limited in Holocene times, but there are cliffs on exposed promontories and island coasts. Sea level rise after the latest glacial period was rapid from −130 m to about the present level, which reached between 7,000 and 5,000 bp. Since that time a wedge of Holocene, mainly terrigenous sediment was deposited, with seaward decreasing thickness. In Gamagyang Bay on the southwest coast the Holocene deposition rate was more than 1.3 mm/year in areas where tidal currents are weak (and probably were during the Holocene); deeper areas remained off islands and headlands where tidal currents are stronger (more than 1.2 m/s). Off Nakdong river in the southeast a prodelta was formed with foreset and bottomset beds reaching down to the water depth of more than 80 m. The Holocene sediment lie on top of a somewhat reworked, subaerially eroded Pleistocene surface, formed during the glacial low sea level of −130 m. On the islands and promontories of southwest South Korea (>Fig. 20.4.1) there are numerous small beaches of variable composition, ranging from fine sand to coarse gravel and small spits of various shapes and sizes, alternating with rocky promontories. In bays and between the islands, where there is protection against high waves, tidal flats have been formed, increasing in size and occurrence

along the south coast with the increase in tidal range and current velocities. Where large streams flow out, tidal flats have been built up at the river mouth in a distinct deltaic pattern. Tidal flats are sandy where the coast has been eroded or where tidal currents or waves are strong. Shelly banks have formed and migrated shoreward during storms and typhoons. The upper parts of the flats are often covered by periglacial debris (angular pebbles in a fine matrix), which has moved down by solifluction from the adjacent slopes. They have grown upward and outward and are of variable thickness, up to 25 m, over a buried bedrock surface. Salt marshes occur in the upper intertidal zone but have been extensively embanked and reclaimed for agriculture. Where sea walls have been built there has been erosion of the adjacent tidal flats, as at Daeho. On much of the coast the tidal flats are up to 10 km wide, and having no seaward barrier are open to storm waves coming from the Yellow Sea. The tidal flat sediment consist of 80–90% mud and there is a variable morphology with step-like terraces, where the elevation changes several metres over a distance of only a few kilometres. There are no large channel systems, but only locally dense drainage networks. Most channels are ebbdominated and during storm surges the soft fluid mudlike (5–10 cm thick) sediment surface is eroded. The eroded sediment returns with the flood tide during fair weather conditions, with accretion and seaward progression of the flats as the net result. Because of erosion and redeposition the surface of the flats can vary by up to 30 cm in elevation. Landward of the flats usually is a dyke or sea wall and only rarely a salt marsh, most of those that existed have been reclaimed since 1900. Sediment accumulation rates are highest on the middle tidal flats and decrease both seaward and landward. The drainage systems are rather stable with relatively little infilling or formation of channel deposits. The higher parts of the flats are usually strongly or completely bioturbated and locally covered by periglacial debris (angular pebbles and a fine matrix), which has moved downward by mass flow from the adjacent slopes. (Wells et al. 1990; Alexander et al. 1991). Cyclic tidal variations on the west coast are reflected in the varying thickness of sediment laminae deposited mainly on the upper tidal flats, as in Kyunggi Bay (Choi and Park 2000; Choi et al. 2001). The inner flats are mostly muddy and intensely bioturbated, the middle flats are more sandy and wave-bedded and the outer flats sandy and ripple-laminated. The Holocene tidal deposits overlie the earlier (interglacial) tidal deposits with sometimes filled-in channels, which are separated from the Holocene by a former subaerial irregular surface, formed before

  20.4

South Korea

⊡⊡ Fig. 20.4.1 Coastal morphology of SW South Korea. (Courtesy Geostudies.) 125°

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12,000 bp. On this surface there is evidence for soil formation and former vegetation (plant roots and fragments). Authigenic microconcretions of siderite indicate nonmarine conditions at about 8,000 bp. Holocene tidal deposition started during the Holocene transgression about 6,000–5,000 bp and gradually reached the present level (it was at −2.5 m in 2,000 bp). This is in contrast to the sea level rise along the Korean south coast, which reached the present level in 5,000 bp at the latest; this is probably related to the slow subsidence of the west coast. The Holocene deposits become gradually coarser upward and seaward and finer inward, which is influenced by the Holocene transgression and the progradation of the tidal flats (Wells et al. 1990; Alexander et al. 1991). Tidal deposits in Gomso Bay are sandier, and an 800-m long shelly sand ridge (chenier) has been formed as early as 1,800 bp high on the flats. It is being built up by broad sand shoals on the middle flat, which during the past 20 years has moved landward. The ridge itself is

also moving landward at an average of 8 m/year. During storms the displacement is two to three times greater: during one typhoon the ridge was displaced over 11 m in a few days. There are some sandy beaches that are backed by dunes, which are up to 30–40 m high near Manripo. They are aligned N-S or NW-SE with the prevailing wind direction. Most dunes are covered by vegetation, but mobile barchans occur locally. Below and behind the recent dunes, which still receive sand from nearby beaches, there are often older dunes of yellow brown sand that probably date from the early Holocene. Off the west coast shallow submerged sand ridges, up to 80 km long, about 2.5 km wide and with an amplitude of up to 30 m, have been shaped by the strong tidal currents. Similar sand ridges probably formed between 10,000 and 7,000 bp, during a period of slow sea level rise, and are now on the sea floor at a depth of between 50 and 90 m.

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South Korea

⊡⊡ Fig. 20.4.2 Coastal geology of Yeongil Bay. (Courtesy Geostudies.)

On the south coast hilly promontories and islands are separated by irregular bays. The tide range is smaller and mudflats and salt marshes less extensive. Pusan stands on a mountainous peninsula to the east of a bay into which the Naktong River has built a substantial delta. Offshore is the island of Cheju-do, a recently extinct volcanic mountain rising to 1,950 m with a coast of low cliffs cut into lava and ash.

The east coast of South Korea is much simpler in outline than the west and the south coasts and is shaped by NNE swell. Between Ulsan and Pohang the coast is steep and rocky, with well-developed terraces up to 130 m above sea level. Six levels have been distinguished, the lowest being the only locally developed and sedimentcovered beach at 2–7 m. Peat in sediment at 2 m has been dated at about 1,700 bp, whereas a terrace at 13 m has

  20.4

South Korea

been dated from charcoal at about 12,000 bp. This gives an uplift during the Late Quaternary of 1.1–1.4 mm/year, but alluvial deposits along the east coast indicate that sea level reached its present level there around 6,000 years ago, implying that no uplift occurred since that time. Near Ulsan is the drowned valley of the Dong River, forming a large estuary. To the north a peninsula points northeastward into the Sea of Japan at the northern end of a low coastal range which is separated from the mainland mountain ranges by a broad graben. At the northern end of this graben a large bay has been formed, Yeongil Man, with the Hyungsan iver flowing into it largely closed off with beach ridges which curve northwest to the city of Pohang (> Fig. 20.4.2). To the north a narrow coastal plain borders the Taebaek Mountains. Small rivers, coming down from the mountains, flow out across the alluvial plain with spits and beach ridges and through narrow gaps to the sea. Further north the coast steepens and emerged terraces become prominent, parallel to the present coastline. North of Geojin the coast consists of large alluvial plains and river mouths with beach ridges and spits shutting off lakes or

lagoons from the sea. Between the plains there are rocky capes and headlands.

References Alexander CR, Nittrouer CA, DeMaster DJ, Park YA, Park SC (1991) Macrotidal mudflats of the southwestern Korean coast: a model for interpretation of intertidal deposits. J Sediment Petrol 61:805–824 Choi KS, Park YA (2000) Late Pleistocene silty tidal rhythmites in the macrotidal flat between Youngjong and Yongyou Islands, west coast of Korea. Marine Geology 167:231–241 Choi KS, Kim BO, Park YA (2001) Late Pleistocene tidal rhythmites in Kyunggi Bay, west coast of Korea: a comparison with simulated rhythmites based on modern tides and Implications for intertidal positioning. J Sed Research 71:680–691 Guilcher A (1976) The ria coasts of Korea and their morphological evolution (in French). Ann Geogr 85(472):641–671 Park YA (1969) Submergence of the Yellow Sea coast of Korea and stratigraphy of the Sinpyeongcheon marsh, Kimje, Korea. J Geol Soc Jpn 5(1):57–66 Wells JT, Adams CE, Park YA, Frankenberg EW (1990) Morphology, sedimentology and tidal channel processes on a high tide range mudflat, West coast of South Korea. Mar Geol 95:111–130

1211

20.5  Japan

Kazuyuki Koike*

1.  Introduction The Japanese islands are located in the northwestern Pacific, from 24°N to 45° 30’ N. The climate is temperate oceanic, with warm summers and mild winters over much of the country. However, in the northwest, Honshu and Hokkaido have cold winters when the sea freezes, and springs when the beaches are covered with blocks of drifted sea ice, while the southernmost islands are subtropical, fringed by coral reefs or mangrove swamps. Tokyo has mean monthly temperatures of 4.7°C in January, rising to 23.2°C in July, with an average annual rainfall of 1,460 mm, while Sapporo in Hokkaido has mean monthly temperatures of −4.9°C in January and 20.2°C in July, with an average annual rainfall of 1,158 mm. The coastline of Japan (> Fig. 20.5.1) is about 34,000 km long, including the coasts fringing the smaller islands. The coastlines of the four major islands total about 18,000 km, one-sixth of which is sandy beaches or deltas, the rest being cliffed, with or without narrow beaches or pocket beaches. Most of the coastal lowlands are now so intensively used that a large ­proportion of the Japanese coast is man-made. A survey by the Environmental Agency of Japan revealed that only 49% of the major island coasts remained natural, 15.6% being semi-natural (occupied by roads, shore protection, and other artificial structures, but still natural in the intertidal zone), and 34.1% consisting of artificial coasts, with marked changes caused by port construction, reclamation, and other developments. The Japanese mountains have been so greatly dissected by torrential rivers that the contemporary rate of fluvial sediment supply is very high, reaching a maximum value of 9,817 m3/sq km/year in the upper catchment of the Kurobe in central Japan. The sediment supply from rivers has been sufficient to maintain accretion of sandy and deltaic coastlines in Japan, and human activities have accelerated coastal erosion during the last few decades, especially near the mouths of the rivers (Uda 1997). A reduction of fluvial sediment supply has caused coastline retreat on a large scale. The coastal landforms of Japan are varied. There are ria coasts and gently-curved sandy coastlines, and it is

common for steep or cliffed coasts to be fringed by a flight of Quaternary marine terraces that have been uplifted, tilted, and faulted in many places. There are also volcanic sectors, formed by eruptions close to the sea. Numerous references to rocky coasts in Japan are listed by Sunamura (1992). The mountainous coasts of Japan were submerged and became deeply embayed during the Postglacial (Jomon or Holocene) marine transgression, which culminated about 6,000 years ago. Many embayments have been filled with sediment supplied by rivers or from nearby sea cliff erosion to form sandy depositional coasts, but some embayments remain unfilled because they have not received sediment from large rivers. Wave action on the Pacific coast of Japan includes ocean swell and occasionally larger waves, with heights of up to 9 m and periods of 14–18 s, generated by typhoons in summer (usually in September). On the Sea of Japan coast, the strongest waves are those generated by the northwesterly monsoon in winter, with heights of up to 5 m and periods of 8–10 s. Patterns of longshore drifting have been analysed by Mizumura et al. (2000) on the Kotogahama coast. Mean spring tide ranges are up to 2.4 m on the Pacific coastline, 0.5 m on the Okhotsk coast, and only 0.2 m on the Sea of Japan coast. Yokohama has a mean spring tide range of 1.4 m and Kamaisi 1.0 m. Higher ranges are recorded around the Seto Inland Sea (up to 2.7 m) and Ariake Bay in Kyushu (up to 4.9 m).

2.  The Japanese Coastline On Hokkaido sandy coastlines are well developed, and most still remain natural. Beaches are covered and eroded by blocks of drift ice in winter along the coasts of Okhotsk and eastern Hokkaido. Several kinds of beach topography are formed by sands drifting along the coast. In this area, embayments formed by the postglacial marine transgression have remained as lagoons behind spits or barrier islands built across bay mouths, as at Saroma, Tokoro, and Yufutsu. Notsuke-zaki is the best developed compound recurved spit  in Japan, and a spit has formed at the mouth of the

*Edited version of chapter 20.5 (Japan) in The World’s Coasts: Online (2003) by Kazuyuki Koike (Komazawa University, Japan). All Rights Reserved. Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_20.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

20.5

Japan

⊡⊡Fig. 20.5.1 Predominant geology of the Japanese Islands. 1– Quaternary sediment 2 – Neogene sediment 3 – Palaeogene and older rocks; 4 – Quaternary volcanics. Nearshore topography: 1 – beaches with smooth profiles; 1a – steep; 1b – gentle; 2 – beaches with stepped profiles; 3 – beaches with submarine bars; 3a – single bar; 3b – multiple bars. Mean spring tide ranges indicated in metres. 0.2

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Japan

20.5

⊡⊡Fig. 20.5.2 Spit formed near the mouth of the Syunbetsu River, Nemuro Strait, Hokkaido.

⊡⊡Fig. 20.5.3 Receding cliff cut in Pliocene formations capped by marine terrace gravel on the Iwaki coast, Hukushima Prefecture.

Syunbetsu River, bordering Nemuro Strait (>Fig. 20.5.2). Northeast of Hokkaido the Kuril Islands are a chain of volcanoes, several linked by tombolos of sand and gravel. At some places on the coast of Hokkaido volcanoes have erupted, forming volcanic cliffs, but much of the Hokkaido coast has outcrops of Tertiary and older rocks. Cliffs on the Hidaka coast are cut in poorly consolidated Neogene sediment. Emerged marine terraces occur on various parts of the coast, notably on the Oshima Peninsula. Southward drifting of beach material on the coast of Tartar Strait has produced a series of asymmetrical spits, some with hooked terminations. Tsunamis have damaged structures on the Hokkaido coast, as on the southern tip of Okushiri Island and on Kenbokki Island.

Most of the coast on the Pacific coast of northeastern Honshu is cliffed. Sanriku has a typical ria coast, which has frequently suffered violent tsunamis associated with giant earthquakes, mainly centred in the Japan Trench. Maximum inundation was up to 28.7 m in Ryori Bay in 1933. Breakwaters have been constructed across the mouth of Ofunado Bay to protect industrial and residential areas from such tsunamis. Matsushima Bay, near Sendai, is famous for its small islands covered with black pine trees. Some of the small islands have been incorporated into the straight sandy coastline, which has been prograding during the past century. The Iwaki coast, south of Matsushima, faces the Pacific Ocean with cliffs 30–35 m high (>Fig.  20.5.3)

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⊡⊡Fig. 20.5.4 Tetrapods at Atami. (Courtesy Geostudies.)

⊡⊡Fig. 20.5.5 Breakwaters for tsunami protection across the entrance to Ofunato Bay, Sanriku Coast, a ria coast subject to tsunamis in Iwate Prefecture.

cut into Pliocene and Pleistocene sediment, which have retreated at a rate of up to 0.7 m/year. Three power plants were constructed along this coast, with ports formed by building breakwaters out from the cliffs. Sand drifting northward along this coast has been partially intercepted by these breakwaters, so that erosion downdrift is due partly to the interception of drifting sand, but also to a decline in sediment yield from a nearby river. Extensive sectors of the Japanese coastline have been armoured with concrete tetrapods to counter beach and

coast erosion, as at Enoshima on Sagami Bay. Atami, on the west coast of Sagami Bay, has a beach completely covered by tetrapods, deposited to protect the seafront from typhoon surges (>Fig.  20.5.4). Breakwaters have been built to protect the coast against tsunamis in Ofuntao Bay (>Fig. 20.5.5). Stacks of Cretaceous rocks border the coast of Ofunato Bay (>Fig. 20.5.6), and cliffs fronted by ironrich sandy beaches border the Shimokita Peninsula in Aomori Prefecture. In Suruga Bay, west of the Shimoda Peninsula, strong wave action arriving through the deep water of the Suruga

Japan

20.5

⊡⊡Fig. 20.5.6 A stack with caves is cut in rocks of the Cretaceous Ofunato Group at Anadori-iso, at the mouth of Ohunato Bay, and is a natural monument of Japan.

⊡⊡Fig. 20.5.7 Notched cliff with shore bench, south of Nabeta Bay.

Trough has produced steep reflective gravel beaches and depositional barriers. Tectonic movements are active here. On the Shimoda Peninsula ocean swell generates strong surf on beaches, as at Nabeta Bay. Cliffs to the south have been notched by strong wave action, with basal high tide shore benches (>Fig. 20.5.7). Rugged cliffs in volcanic rock border the southern end of the Shimoda Peninsula. Sandy coastlines are extensive along the Kanto Plain, which is composed of Pleistocene uplands and Holocene lowlands. The coast around Tokyo Bay has been so highly developed that almost all the deltaic sectors have been

reclaimed and converted to ports and industrial and residential areas. A natural delta coastline and tidal flat persist at the mouth of the Obitsu River. On the west coast, at Kemigawa, the beach is backed by the Inage Seaside Park (>Fig. 20.5.8). South of Chiba an artificial beach has been inserted in front of a reclaimed area, and is intensively used (Koike 1990). A significant progradation of the sandy coastline has occurred in the central part of Kujukurihama, Chiba Prefecture, where the beach advanced at least 200 m between 1903 and 1967. A comparison of an historical

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⊡⊡Fig. 20.5.8 Kemigawa Beach and Inage Sea Side Park on the western coast of Tokyo Bay, Chiba Prefecture. Artificial beaches were constructed in front of the reclaimed area (a former tidal flat) during the 1970s. (Courtesy of the Chiba Prefectural Government.)

⊡⊡Fig. 20.5.9 Byobugaura cliff, northeastern end of Kujukurihama, Chiba Prefecture, in 1976. The cliff is cut in Pliocene rocks, and receded at 1–2 m/year before the construction of protective works. Sand eroded from this cliff drifted southwest to Kujukurihama beach.

(1735) sketch map with recent topographic maps shows that beach accretion of about 400 m has occurred near the mouth of the Sakuta River, in central Kujukurihama. There is fairly good agreement between the area of multiple bar development and that of coastline accretion exceeding 2  m/year. Long-term longshore drifting here is towards the centre of Byobugaura Bay, as inferred from grain size and heavy mineral distributions. The outcome of this convergence is a prograded beach, backed by grassy dunes, at Katakai. North of Kujukurihama is the Byobugaura Peninsula, where the Iioka bluffs, a former coastline, approach the

present coast. The tall cliffs have been armoured at the base, and are gradually degrading (>Fig. 20.5.9). Raised (emerged) Holocene marine terraces are well developed on both the Boso and the Miura peninsulas (>Fig. 20.5.10) and at the northern head of Sagami Bay. These are often referred to as the Numa terrace, after the type locality of fossil coral beds at Numa, Tateyama City. Successive uplifts in the Holocene have produced a flight of marine terraces in front of the Oiso Hills and on the southern part of the Boso Peninsula. Near Chikura, Chiba Prefecture, there are four marine terraces, with altitudes of 25, 16, 12, and 5 m above sea level. The first terrace is

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⊡⊡Fig. 20.5.10 Emerged shore platform on weathered tuff and agglomerate at Misaki.

mainly a depositional surface of marine sediment dated about 6,000 years ago, which is correlated with the fossil coral beds at Numa. The fourth terrace emerged at the time of the great earthquake of 1703. At Kashima an industrial port was constructed on the sandy coast facing the Pacific Ocean. Long breakwaters were built out into the ocean, followed by a Y-shaped inner basin with several kinds of port facilities in the coastal lowland. In Tokai District rivers that drain the Kiso and Akaishi ranges (such as the Fuji, Abe, Oi, and Tenryu) deliver much sediment to their mouths. Their gravelly deltas prograded until the late 1940s, when erosion became prevalent. This is thought to be related to a reduced sediment yield caused by the sediment interception in man-made reservoirs and the extraction of gravels from river beds. On the west coast of Izu Peninsula are gravelly beaches and dissected stacks, and on the south coast deep marine inlets (rias) penetrate submerged valleys in volcanic formations. Volcanic features include the Omuro-yama cone (>Fig. 20.5.11) and lava flows forming benches on Oshima Island. Maars on Miyake-jima Island, near Tokyo, were enlarged by the 1984 eruption (>Fig. 20.5.12). On the Pacific coast of southwestern Honshu a flight of Pleistocene and Holocene marine terraces is extensive, especially along the coasts of the peninsulas bordering the Nankai Trough, such as Kii, Muroto, and Ashizuri. These terraces have emerged and been tilted northward by local tectonic movements. In contrast, a ria coast is found on the northern part of each peninsula. This is attributed to

crustal deformations associated with the giant earthquakes that have occurred along the Nankai Trough. On the western coast of Honshu earthquakes have disrupted landforms and damaged structures on Awaji Island and in Kobe. Cliffs have receded, leaving lines of stacks on the southern tip of Kii Peninsula. The Seto Inland Sea is studded with numerous small islands that are composed mainly of deeply weathered granites. The beaches, consisting of arkose (feldspathic) sands, have been used for salt production, making the most of a relatively large tide range and dry weather in this part of Japan, but they are now being converted to industrial use. Coastal sand dunes are extensive on the coasts bordering the Sea of Japan, where the strong northwesterly winter monsoon blows onshore (Endo 1986). In some parts, wind-blown sands now cover Pleistocene uplands, forming dunes up to 80 m high. Most of the coastal dunes along the Sea of Japan coasts were not covered with vegetation, and were therefore blown inland in the medieval era. Many feudal lords (Daimyo) tried to fix the mobile dunes by planting black pine trees (Pinus thunbergii) to protect paddy fields from inundation by blown sand during the Edo era. Today these dunes are covered with beautiful forests, which have fixed the mobile dunes and trapped sand blown from the shores. Sandy coastlines have retreated markedly in many sectors bordering the Sea of Japan, for various reasons. Near Kanazawa City the coast has tended to retreat in the last 60 years, and rapid erosion would have continued if

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⊡⊡Fig. 20.5.11 Omuro-yama volcanic cone and lava flow, eastern coast of Izu Peninsula, Shizuoka Prefecture.

⊡⊡Fig. 20.5.12 Maars enlarged by the 1984 eruption of Miyake-jima Island, Tokyo.

protection works had not been emplaced on the coast of  Toyama Bay. Rivers draining Toyama Bay, such as the  Kurobe, Joganji, and Jinzu, have steep channels, and supply a large amount of sediment brought down from the mountains. Nevertheless, the coast has been eroded because each river channel flows directly into the head of a submarine canyon, so that a large proportion of debris supplied from the rivers is carried down to suboceanic abyssal plains instead of being distributed along the coast. On the north coast of the Sea of Japan there are cliffs cut in partly consolidated Tertiary formations. There are

emerged shore platforms produced during earthquakes, as at Senjo-jiki. Volcanic features include coastal maars and calderas (>Fig. 20.5.13), and there are extensive dunes at Shonai. Coastline erosion around Niigata City has accelerated since the 1930s, primarily because sand supply to the old mouth of the Shinano decreased after the construction of a diversion channel in 1922, for the purpose of flood control. Sediment yield from the Shinano is estimated at 12 million tons/year, three-quarters of which now flows directly into the sea through the new mouth, where wide

Japan

20.5

⊡⊡Fig. 20.5.13 Ninome-gata maar and bay in Toga caldera, Oga Peninsula, Akita Prefecture, north coast of the Sea of Japan.

⊡⊡Fig. 20.5.14 A cliff 20 m high cut in the Tertiary Aikawa formation and a rocky marine terrace, backed by terraces fronting mountains in Senkaku Bay, western Sado Island.

sandy forelands have been added on both sides of the river mouth since 1922, forming a cuspate delta. Changes here have also been influenced by local land subsidence, largely due to groundwater extraction. An undersea reef has been built to diminish incident waves on the eroding coastline. Rocky coasts on the west of Sado Island, facing the Sea of Japan, are fringed by several well-developed marine ­terraces (>Fig.  20.5.14), which are useful references for the estimation of tilting and faulting as well as vertical

movement in the late Quaternary. Shore platforms show scoured potholes, and there is a tombolo linking Futatsugame at the northern end of Sado Island (>Fig. 20.5.15). The Noto Peninsula is remarkable for rice paddy terraces on coastal slopes (>Fig.  20.5.16). Volcanic coasts are found around Kyushu and Hokkaido. Sakurajima volcano, in Kagoshima Bay, is one of the most active in Japan. It used to be separated from the mainland, but was tied to the Osumi Peninsula by  lava flows in 1914–1915, forming a

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⊡⊡Fig. 20.5.15 Tombolo between main island of Sado and Futatsu-game at the northern tip of Sado Island, Niigata Prefecture. It was damaged by a tsunami in 1983.

⊡⊡Fig. 20.5.16 Man-made terraces for the Sen-maida rice paddy terraces, Noto Peninsula, Ishikawa Prefecture.

natural harbour. There  are emerged fringing reefs on the northeastern coast of Kikaijima and around Yoron Island (>Fig. 20.5.17), and beach rock on the Okinoerabu shore. The Amano-Hashidate spit in Wasaka Bay (>Fig. 20.5.18) has been ­artificially preserved. Sandy coasts are well preserved along the Pacific side of Kyushu. There are shore platforms, tombolos, and river-mouth spits around Aoshima, Miyazaki Prefecture (>Fig. 20.5.19). Submergence has resulted in the formation of rias. Most of the western coast of Kyushu facing the

East China Sea is cliffed, and in some parts volcanoes stand directly beside the sea. Deltaic coastlines are prograding around Ariake Bay, which has the highest tidal range (up to 5 m) in Japan. Two island arcs extend southward into the Pacific Ocean, the Izu-Ogasawara arc from Kanto (Kaizuka et al. 2000), and the Ryukyu arc from Kyushu, each reaching as  far as 24° N. These are double arcs, with outer ­non-volcanic islands and inner volcanic islands. There are  emerged coastlines in the Ryukyu islands. Coral

Japan

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⊡⊡Fig. 20.5.17 A fringing reef around Yoron Island.

⊡⊡Fig. 20.5.18 The Amano-Hashidate sand spit in Wakasa Bay, Kyoto Prefecture, Sea of Japan. This has been a famous scene for at least a thousand years. The spit was eroded after the 1950s, and is now protected by a series of small groynes and annual renourishment with 4,800 cubic metres of sand.

reefs  (some emerged) now fringe most of the Ryukyu Islands and there is a mangrove swamp on Iriomote-jima, southernmost of the Ryukyu Islands. At  Agari Hen’nna Cape, on the east coast of Mikayo Island in the southern Ryukyu  Islands, boulders strewn across a fringing coral reef are thought to have been emplaced by a ­tsunami. Erosional lowering of the shore platform has produced notches and mushroom rocks on Okinawa Island (>Fig. 20.5.20). Nishino-shima Shinto is a volcanic island located in the southern part of the Izu-Ogasawara Arc that began to

form soon after a submarine eruption occurred, about 800 km south of Honshu, in April 1973. Surveys by the Japanese Marine Safety Agency showed that the new island continued to grow by successive eruptions until August 1974. The tendency thereafter has been towards a reduction in the size of the island as volcanic activity subsided and coastal erosion prevailed. An abrasion platform, more than 100 m wide, has formed off the southern coast of the new island, extending to about 15 m below sea level, the depth to which vigorous wave abrasion occurs in this typhoon-prone environment.

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Japan

⊡⊡Fig. 20.5.19 Shore platform south of Aoshima, Miyazaki Prefecture, Kyushu.

⊡⊡Fig. 20.5.20 Emerged notches and mushroom rocks above a reef flat on the southern tip of Okinawa Island.

Southeast of Japan, Iwo-jima is a rapidly rising volcanic island (Kaizuka 1992).

References Endo K (1986) Coastal sand dunes in Japan. Proceedings of the Institute of Natural Scienes, Nihon University, 21:37–54 Kaizuka S (1992) Coastal evolution at a rapidly uplifting volcanic island: Iwo-jima, western Pacific Ocean. Quat Int 15/16:7–16

Kaizuka S, Koike K, Endo K, Yamazaki H, Suzuki T (eds) (2000) Geomor­ phology of Kanto and Izu-Ogasawara. Regional ­Geomorphology of the Japanese Islands, Vol 4. University of Tokyo Press, Tokyo, Japan Koike K (1990): Artificial beach construction on the shores of Tokyo Bay. J Coastal Res Special Issue 6:45–54 Mizumura K, Tagawa K, Yamamoto T (2000) Seasonal changes of direction of sand movements on Kotogahama coast. Env. Geol 39:641–648 Sunamura T (1992) Geomorphology of rocky coasts. Wiley, Chichester Uda T (1997) Coastal Erosion of Japan. Sankaido, Tokyo

21.0 Australia – Editorial Introduction

The Australian continent has an area of 7.68 million sq km. It is situated on a continental plate, and has been raised by post-Cretaceous epeirogenic movements, acc­ ompanied by tectonic uplift and folding of the Eastern Highlands (Jenkin 1984). Tectonic activity is now rare, and stability accompanied by long-continued planation has produced the flattest of the continents, the highest summit, Mount Kosciusko, being only 2,228 m above sea level (Gill 1982). Partly because of its subdued relief, Australia is also the driest continent, and fluvial water and sediment yields to the coast are small (>Fig. 21.0.1). It is noteworthy that the largest river system, draining the Murray-Darling Basin (1.06 million sq km) has a mean annual runoff of only 15,000 m3 per sq km, about 1.25% of that of the Yangtze-Kiang in China. Depleted by aridity (and more recently by irrigation use), the Murray flows into coastal lagoons, and delivers no sediment to the coast. There is also much internal drainage, notably in the Lake Eyre Basin (1.17 million sq km), which makes no contribution to the coast. Two-thirds of the continent is arid, but there are humid temperate zones in the southwest and southeast (including Tasmania) and humid tropical zones in the northeast and north (Radok 1976). Maximum runoff is in the winter in the south and summer in the north. The length of the Australian coastline (mainland  + Tasmania), based on the counting of 1 km intercepts, is about 26,900 km (Galloway and Barr 1979), increasing to about 41,000 km if the coastlines of smaller islands (area > 12 ha) are included. Wave action is dominated by swell from the Southern Ocean, with high wave energy arriving from the southwest along the western and southern coasts and refracted to southerly and southeasterly in Tasmania, eastern Victoria, New South Wales and southern Queensland as far as Fraser Island (>Fig.  21.0.2). Wave regimes are much influenced by the width and gradient of the continental shelf, and waves generated by local winds become important where ocean swell is attenuated or excluded, as in gulfs and bays. On the northwest and north coasts swell is weakened across a wide continental shelf and wave energy is generally low, except during occasional tropical cyclones which generate high winds and

storm surge flooding. The northeast coast is protected by the Great Barrier Reef from Coral Sea swell, and has generally low to moderate wave energy produced by southeasterly trade winds in coastal waters, although here, too there are occasional tropical cyclones with strong winds and storm surge flooding (Davies 1977). Microtidal conditions (mean spring tide range  4.0 m) sector between Darwin and Port Hedland: the maximum spring tide range is 10.5 m in Collier Bay on the Kimberley Coast of northwest Australia (Easton 1970). The large tide range, broad continental shelf and numerous scattered reefs and islands reduce wave energy from the Timor Sea. Much of the western and southern coast of Australia has beaches and dunes of calcareous sand, largely biogenic sediment derived from the adjacent sea floor by onshore drifting by ocean swell during high sea level phases and by wind action during low sea level phases, when the sea floor emerged as a dry coastal plain. Pleistocene dunes have been lithified to dune calcarenite (calcareous aeolianite) on these western and southern coasts. By contrast, beaches and dunes on the southeastern, eastern and northern coasts are generally quartzose, derived from the weath­ering of granites and sandstones in the hinterland and delivered to the coast by rivers, eroded from nearby cliffs or from the adjacent sea floor by onshore drifting of sediment from weathered outcrops of granite or sandstone, or from fluvial or aeolian sands deposited during low sea level phases. They often contain heavy mineral sand horizons. Quartzose sand dunes remain unlithified (apart from the formation of subsoil humates known as coffee rock in the zone of fluctuating water tables), and if their vegetation cover is sparse, or has been reduced by clearing, burning or grazing, they become unstable, with blowouts which may develop into transgressive dunes. On the northeast, north and northwest coasts beaches include coralline sand and gravel derived from fringing and nearshore coral reefs (Bird 1978). Systematic studies of the morphodynamics of Australian beaches (Short and Wright 1984) have been used as the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

21.0

Australia – Editorial Introduction

⊡⊡ Fig. 21.0.1 Australia, showing major river systems. (Courtesy Geostudies.) A

SE

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basis for accounts of beach systems on a state basis, starting with New South Wales (Short 1993). Beach erosion has become a widespread problem in Australia (Bird 1988), attributed to various causes, notably a decline in sediment flow from the sea floor, rivers, and cliffed coasts stabilised by sea walls (which have also intensified wave reflection scour), and a rising sea level. The dominance of shoreward drifting of sand by ocean swell on the western, southern and eastern coasts of Australia and the fact that most rivers discharge into estuaries or lagoons rather than to the coast has resulted in the formation of extensive depositional barriers (Jennings and Bird 1967; Bird 1973). These have beaches shaped by slightly refracted ocean swell, backed by beach and dune ridges (and dune calcarenite ridges) in parallel alignments representing intermittent progradation, sometimes interrupted by blowouts that grow into parabolic dunes and eventually

transgressive sand areas. The Ninety Mile Beach in southeastern Australia borders a series of such barriers. They are typically backed by lagoons occupying former embayments or river-mouth inlets, which are modified by swamp encroachment and the building of deltas by inflowing rivers and re-shaped by waves and currents, some shores eroded, others prograded, so that initially elongated lagoons become segmented into chains of rounded lagoons (Bird 1967). Where longshore drifting of sand becomes more important, as in bays and inlets in southern Australia, and along the Queensland coast behind the Great Barrier Reef, spits have formed, and there are cuspate spits and forelands in zones of drift convergence, notably in the lee of nearshore reefs and islands. Cheniers are ridges of sand and gravel deposited by storm surges on coastal plains (Short 1989). Cliffed coasts are extensive, often with bordering shore platforms. On the western and southern coasts cliffs

21.0

Australia – Editorial Introduction

⊡⊡ Fig. 21.0.2 Some factors influencing the Australian coast. (Courtesy Geostudies.) A

SE R IMO

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have been cut into Pleistocene dune calcarenites, and bordering subhorizontal shore platforms, exposed only at low tide, have been shaped partly by solution and bioerosion as well as by wave abrasion. Vertical cliffs have been cut in Tertiary limestones and sandstones on the coast of the Nullarbor Plain (a plateau) at the head of the Great Australian Bight and on similar rocks in western Victoria, while cliffs cut in Palaeozoic sandstones occur in central New South Wales and northwest Australia. Cliffs have been cut in Tertiary and Quaternary basalts in western Victoria and dolerites in Tasmania. Subhorizontal shore platforms close to mean high tide level on sandstones and volcanic formations are the outcome of weathering (notably recurrent wetting and drying) and the removal of weathered material by wave action. In addition, there are

Bass Strait

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seaward-sloping intertidal shore platforms formed by abrasion, notably where waves mobilise sand and gravel to scour the rock surface. Segments of such abrasion ramps are often minor components of subhorizontal shore platforms. There is more intricate variation in cliffs and shore platforms related to the geological diversity presented by the folded and faulted rock formations of the Eastern Highlands and Tasmania than on the more uniform calcarenite coasts to the west. There is evidence that the sea has stood at various levels relative to the land around the Australian coast during Quaternary times (Hopley 1983). These changes have generally been attributed to eustatic movements (sea level rise and fall), but although Australia has been a generally stable continent tectonic movements have occurred locally,

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Australia – Editorial Introduction

particularly in the southeast of South Australia and along the Victorian coast, tilting and displacing Pleistocene ma­­ rine terraces, raising promontories (such as the Otways coast and Mornington Peninsula) and lowering intervening areas (such as Port Phillip and Western Port Bays). Reference has been made to coral reefs, which extend from Houtman Abrolhos off the west coast round northern Australia to the Great Barrier Reef, ending north of Fraser Island on the east coast (Fairbridge 1967). Man­ groves occupy the upper intertidal zone in estuaries, inlets and sheltered embayments in northern Australia, extending as far south as Bunbury (33° 20ʹ S) on the west coast and Corner Inlet (38° 55ʹ S) in the south east (Clough 1982). In arid sectors they are backed by saline mudflats, but in humid sectors they have forested hinterlands, except where these have been cleared, notably for sugar cane farming. Salt marsh vegetation is found landward of some mangrove areas, and salt marsh takes the place of mangroves beyond their poleward limits, as in Tasmania. In southeastern Australia the European rice grass (Spartina) has been introduced to a number of estuaries, and has become dominant in Anderson’s Inlet in Victoria and the Tamar estuary in Tasmania (Bird 1974). Further information on coastal geomorphology is given in the chapters on the Australian states (> New South Wales, > Queensland, > Northern Territory, > Western Australia, > South Australia, > Victoria, > Tasmania).

References Bird ECF (1967) Coastal lagoons of southeastern Australia. In: Jennings JN, Mabbutt JA (eds) Landform Studies from Australia and New Guinea, pp 365–385 Bird ECF (1973) Australian coastal barriers. In: Schwartz ML (ed) Barrier Islands, pp 410–426

Bird ECF (1978) The nature and source of beach materials on the Australian coast. In Davies JL, Williams MAJ (eds) Landform evolution in Australasia. Australian National University Press, Canberra, pp 144–157 Bird ECF (1974) Assessing man’s impact on coastal environments in Australia. The impact of human activities on coastal zones, U.N.E.S.C.O., pp 66–75 Bird ECF (1988) The future of the beaches. In: Heathcote RL (ed) The Australian experience. Longman Cheshire, Melbourne, pp 163–177 Clough BF (ed) (1982) Mangrove ecosystems in Australia. Australian National University Press, Canberra Davies JL (1977) The coast. In: Jeans DN (ed) The natural environment. Australia, A Geography 1, pp 203–222 Easton AK (1970) The tides of the continent of Australia. Horace Lamb Centre, Research Paper 37 Fairbridge RW (1967) Coral reefs of the Australian region. In: Jennings JN, Mabbutt JA (eds) Landform Studies from Australia and New Guinea, pp 386–417 Galloway RW, Barr ME (1979) What is the length of the Australian coast? Aust Geogr 14:244–247 Gill ED (1982) Eight coasts of Australia. C.S.I.R.O. Division of applied geomechanics, Technical Report 119 Hopley D (1983) Australian sea levels in the last 15,000 years. Occasional Paper 3, Monograph Series, Department of Geography, James Cook University, Townsville Jenkin JJ (1984) Evolution of the Australian coast and continental margin. In: Thom BG (ed) Coastal geomorphology in Australia. Academic Press, Sydney, pp 23–42 Jennings JN, Bird ECF (1967) Regional geomorphological characteristics of some Australian estuaries. In: Lauff GH (ed) Estuaries: American Association for the Advancement of Science, Vol 83. pp 121–128 Radok R (1976) Australia’s coast, an environmental atlas with base-lines. Rigby, Adelaide Short AD (1993) Beaches of the New South Wales coast. Australian Beach Safety and Management Program, Sydney Short AD (1989) Chenier research on the Australian coast. Mar Geol 90:345–351 Short AD, Wright LD (1984) Morphodynamics of high energy beaches: an Australian perspective. In: Thom BG (ed) Coastal geomorphology in Australia. Academic Press, Sydney, pp 43–68

21.1 New South Wales

Bruce Thom

1.  Introduction The New South Wales coast is essentially a partially submerged embayed coast in which the bays are fully or partially filled by sandy bay barriers, tidal flats, lagoons, and deltaic plains (Langford Smith and Thom 1969; Chapmen et al. 1982). Depositional sequences within embayments appear to increase in age, extent, and morphological as well as stratigraphical complexity, from Cape Howe in the far south of New South Wales into southern Queensland. Underlying estuarine mud sequences of late Quaternary age are also more complex in the larger embayments north of New­castle. The nature and size of individual bays and their distance apart is influenced markedly by bedrock type and orientation, and the direction of approach of predominant waves. Many bay barriers have the shape of a half-heart or zeta, producing a distinctive offset coastal outline. Langford-Smith and Thom (1969) divided the coast of New South Wales into two general categories: rugged, where the immediate hinterland is hilly or mountainous, and subdued, where relief is lower, embayments broad, and headlands relatively less conspicuous. In general, rocky cliffs and headlands predominate south of Newcastle, subdued and sandy coasts to the north. The main structural units east of the Great Divide provide the basis for the three coastal provinces discussed below: southern fold belt, central Sydney basin, and northern fold belt and basin provinces (Roy and Thom 1981). Two different rock groups occur along the New South Wales continental margin. The oldest group forms the fold belts: Lachlan in the south and New England in the north. Rocks forming these fold belts were laid down in the Tasman Geosyncline in early to mid- Palaeozoic times. Folding took place in the late Palaeozoic, with associated plutonic intrusions and metamorphism. The youngest group of rocks are of late Palaeozoic to mid-Mesozoic age, and form two large sedimentary basins: the Sydney Basin and the Clarence-Moreton Basin. Horizontal bedding and vertical jointing distinguish the younger group from the strongly deformed rocks of the fold belts. Each group imparts a distinctive signature to the morphology of the headlands and valleys of this coast.

The existing linear continental edge, which cuts across all these rock groups, formed in the late Mesozoic as a result of continental break-up and sea floor spreading. Plate separation that opened the Tasman Sea terminated in the mid-Palaeocene (about 60 million years ago). The western margin of the asymmetrical rift constitutes the East­­ern Highlands of New South Wales, and is bordered by a narrow continental shelf which drops steeply to 4,500  m into the Tasman Sea Basin. Rivers draining the highlands descend to base levels carved across the narrow shelf, which has generally subsided since its formation. The outer part of the shelf comprises a sediment wedge about 500 m thick and 20–30 km wide of mid- to late Cainozoic age. River valleys have progressively become drowned and partially filled in association with sea level rise, in particular as Quaternary high (interglacial) sea levels reached their highest levels along this coast over the last 300,000–400,000 years. Valley extension along a linear margin provides the structural framework for coastal sedimentation in the late Quaternary. Essentially, the coast is bedrock controlled. Marine cliffs, rock platforms, and abrasion surfaces have been carved into the seaward ends of drowned valley systems. Sediment supplied from rivers since Tertiary times to the continental shelf have been extensively sorted and reworked under high-energy wave conditions from the southeast. Sand dominates inner shelf beach dune environments, the sand particles being composed of quartz and other stable minerals, including the economically significant heavy minerals, rutile and zircon. Marine transgressions accompanying deglaciations have swept these sands into embayments as bay barriers or flood-tide deltas, or have piled them against cliff-faces as inner-shelf sand bodies (Roy et al. 1994). Strong onshore winds, especially from the southeast, have transported vast amounts of sand into foredune ridges, transgressive dune complexes, and cliff-top dunes. As sea level has been close to its present position over the past 6,500 years, the stillstand following the end of the Holocene marine transgression, estuaries have become totally or partially filled with quartzose from the sea floor and from terrestrial sources (gravels, sands, and muds). Bay barriers have reformed during the Holocene,

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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­ igrating onshore during the marine transgression and m consolidating during the ensuing stillstand. In summary, seven key events constitute the geological framework of the New South Wales coast. There has been Palaeozoic formation of fold belts, including granitic intrusions; Permo-Triassic sedimentary basin deposition; the Palaeocene opening of Tasman Sea, forming the continental margin; Tertiary river incision into the Eastern Highlands and across the shelf edge; Quaternary subsidence of the continental shelf with continued deposition of a sediment wedge and the carving of cliffs; Quaternary oscillations of sea level, with active shelf sand transport, sand barrier deposition, and valley fill during high interglacial sea levels; and the Late Pleistocene to Holocene marine transgression and stillstand, with valley drowning, onshore sand transport, and estuary infilling. The New South Wales coast is oriented NNE–SSW, and is subject to a dominant south to southeast swell and wind-driven wave regime generated in the south Tasman Sea. Due to the narrow and relatively deep continental shelf, frictional loss of wave energy is negligible until close to the surf zone. As a result, the coast is dominated by a high-energy wave climate. Deep water waves up to 16 m high have been recorded. Incident wave energy on the central New South Wales coast is 5–10 times greater than on the Atlantic coast of North America, and 80 times greater than on the Gulf of Mexico coast near the Mississippi delta. Because of the relationship between dominant wave climate and coastal orientation, there is a tendency towards northward longshore drifting. However, in the southern part of the state (south of 32° S), embayments are compartmentalised either by prominent headlands that extend into deep water, or by deep estuary mouths. Barriers to longshore drifting are less frequent to the north, and by-passing occurs between most embayments. Longshore drift rates of as much as 500,000 m3/ year occur in southern Queensland, which is an extension of the New South Wales sediment transport system. Coastal climatic conditions in New South Wales are characterised by variability. Five meteorological systems influence rainfall, temperature, and oceanic conditions: tropical cyclones (especially north of 32° S); mid-latitude cyclones (travelling west to east); east-coast lows; zonal anticyclonic highs; and summer sea breezes. Annual rainfall along the coast is between 1,800 and 1,300 mm. Storm waves may be generated at any time of the year, but there is some seasonality, with the months of February, March, and June experiencing the largest wave heights. However, even in September, one of the least stormy months, devastating storms (such as the one in 1974) can occur, inducing massive coastal erosion. The variable incidence of storms

enables beach and foredune morphologies to respond ­episodically, with lengthy periods of accretion followed by  shorter periods of erosion (McLean and Shen 2006). Climatic variability is also expressed within embayments by beach rotation (Short et al. 2000). El Niño and Pacific Decadal Oscillation effects can be inferred from detailed beach surveys at a range of sites, indicating the importance of long-term monitoring in order to understand the variable dynamics of New South Wales beach systems. Tide records also reveal patterns of oscillation that have signals extending over decades, as well as interannual variations. The Fort Denison gauge in Sydney Harbour provides information extending back to 1894. sea level appears to have risen by about 0.5–1 mm/year over this period at this site. Storm surges and even minor tsunami events are evident in these records, where the maximum astronomical tidal range is about 2 m. More recent works have revealed high sea level anomalies of the order of 20–30 cm, unrelated to weather phenomena, that help to induce beach erosion. The impact of the south-flowing East Australian Current with its eddies and temperature patterns is relatively little understood in relation to sea level change, meteorology, climate change, biology, and even sediment transport.

2.  Southern Province The southern province lies within the Lachlan Fold Belt. It  stretches from the Victorian border to just north of Batemans Bay, a funnel-shaped estuary at the mouth of the Clyde River. This province is characterised by a vast array of creeks and rivers draining relatively small, forested catchments. The valleys cut across a range of N–S striking rocks, mostly of Ordovician to Devonian age. Emplacement of late Palaeozoic intrusive rocks ensures that coastal scenery is locally dominated by massive plutonic rocks. Weathering of such rocks inland yields coarse, feldspathic-rich sands, which, at a few locations, reach the estuary (as on the Moruya River) or even the beach: Whale Beach in Twofold Bay is fed by the Towamba River (>Fig. 21.1.1). Tertiary basalts, often found in association with fluvial gravels (>Fig. 21.1.2), also occur on stretches of this coast, adding to its geological diversity. Typically, embayments are narrow, between steepsided ridges terminated at their seaward ends by cliffs. Shore platforms fringe these cliffs and truncate the intricate fold patterns of the Palaeozoic sediment and metasediments (Bird and Dent 1967), but there also structural benches that coincide with the upper surfaces of resistant flat or gently sloping formations, prominent in

New South Wales

21.1

⊡⊡ Fig. 21.1.1 Whale Beach, at the mouth of Towamba River, consists of coarse quartz sand with abundant feldspar supplied by the river. It contrasts with most New South Wales beaches, which are not fluvially nourished and contain little feldspar. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.2 The shingle beach at Kiama, south coast New South Wales, formed from disintegrating basalt. (Courtesy Geostudies.)

some cliff profiles (>Fig. 21.1.3). Emerged benches occur locally, formed during phases of higher relative sea level. Dissection of cliff outcrops yields coves, caves, and blowholes (>Fig. 21.1.4). Subhorizontal shore platforms are well developed on fine-grained sandstones, mudstones, and basalt outcrops (>Fig. 21.1.5), formed by weathering processes, including recurrent wetting and drying, which reduces the shore

outcrop to the level of permanent saturation (>Fig. 21.1.6). Below the low tide line the smoothness of platform surfaces is replaced by a sharp micro-relief of ridges and gutters, and even caves and tunnels where the basalt outcrops, as on the south side of Broulee Headland. The reasonably common occurrence of ramparts on the most exposed sections of platforms, apparently irrespective of rock type, is hard to explain.

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⊡⊡ Fig. 21.1.3 Structural benches on Bay Cliff at Wonboyn, south coast New South Wales. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.4 The blowhole at Kiama. (Courtesy Geostudies.)

Beaches occupy embayments between rocky headlands, and are often asymmetrical (zeta-curved) in outline, shaped largely by the dominant southeasterly ocean swell, refracted as it approaches the coast. There are cycles of cut and fill, with erosion during stormy periods and restoration of beach profiles in calmer weather. Spits form in the lee of islands and reefs (>Fig. 21.1.7). Detailed studies have been conducted on the stratigraphy, morphology, and age structure of Holocene marine sand deposits within embayments on the New South Wales coast. Several types of open-ocean sand accumulation have been recognised: prograded barriers involving marine transgression to about 6,000 years ago, then beach ridge and foredune accretion; stationary barriers involving the more or less vertical growth of bay barriers from depths of 10–20 m below present sea level to the present; receded barriers involving the construction of narrow bay barriers about 5,000–6,000 years bp, followed by recession over paludal sediment to the present; episodic dune development involving phases of dune instability over the last 6,000–8,000 years, with mobile sand sheets migrating downwind over vegetated sand, lagoonal, and bedrock surfaces; mainland beaches involving the accumulation of sand at the heads of embayments against bedrock slopes; and inner-shelf sand bodies. Within the southern province there are a number of  sites where drilling has revealed the sedimentary sequence in different bay barriers. Radiocarbon dating of shell hash within the sands has yielded age structures with more or less consistent signatures, although higher

New South Wales

21.1

⊡⊡ Fig. 21.1.5 Shore platform on Broulee Island, south coast New South Wales. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.6 Weathering features on shore platform at Broulee Island. (Courtesy Geostudies.)

resolution dating could reveal anomalies given the mixture of materials that make up a barrier. One such site is the prograded ridge sequence north of the Moruya River (Thom 1984). Within the prograded barrier three distinct sediment units have been defined: a basal estuarine mud, an overlying shelly sand formed during the Holocene marine transgression, and the prograded sand sheet. On promontories at the head of the embayment the sand sheet abuts relic cliffs and buries fringing rock

platforms, which may have formed during previous high sea levels. Within this and other embayments in the southern province there is a paucity of evidence for high interglacial sea level deposits compared to the northern province. Four basic estuary types occur in New South Wales (Roy et al. 2001): drowned river valley, barrier estuary, saline coastal lagoon, and mature river estuary. All four types occur in the southern province. They have been

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New South Wales

⊡⊡ Fig. 21.1.7 A tombolo at Broulee Island. (Courtesy Geostudies.)

studied from both a geologic-geomorphic as well as a biological perspective, and represent a diverse array of sedimentary environments and habitats. Batemans Bay at the north end of the province is a classic drowned valley (Clyde River), with extensive flood-tide shoals extending across its mouth. Drilling below fringing sand beaches reveals a deep (less than 20 m) sequence of estuarine mud, which accumulated in part during the Holocene marine transgression. Modern fluvial sands enter the sequence from the west. Barrier estuaries and saline coastal lagoons may exist side-by-side (e.g. Tuross and Coila Lakes). The key difference is that barrier estuaries have sufficient impact of freshwater (Tuross River), combined with tidal currents, to maintain an open tidal inlet, although closure may occur at times of severe drought. Saline lagoons are often more frequently closed by a beach berm than open, and hence they have been given the acronym of ICOLL (Intermittently Closed or Open Lake or Lagoon). Small coastal creeks can also be intermittent in draining to the sea (Roy et al. 2001). Lakes and creeks open during heavy rain and high energy wave events. Mature river estuaries occur where larger rivers have mostly filled the lake basin behind a barrier. The Bega River estuary is a good example, where fluvial sands and muds have aggraded over tidal flat and estuarine mud-basin deposits to create a deltaic plan within an ancient bedrock valley.

3.  Central Sydney Basin Province North from Batemans Bay to the mouth of the Hunter River at Newcastle, the New South Wales coast is dominated by sedimentary rocks of the Sydney basin. The rocks form a layered sequence of sandstones and shales, with coal measures, conglomerates, and a variety of extrusive volcanics interbedded within the sequence. The oldest rocks of Permian age outcrop at the northern and southern ends of the basin. The thickest sequence (about 1,000 m) occurs in the vicinity of Sydney Harbour, where Triassic rocks overlie those of Permian age. Here, the Hawkesbury Sandstone Formation forms majestic cliffs. Coastal scenery is mostly fashioned out of massive and interbedded sandstone units, rich in quartz; the sandstone is typically horizontally bedded and vertically jointed. As in the Southern Province, ancient river valleys carve through the basin rocks from the Eastern Highlands to the west. Deeply incised, these valleys extend across apparently planed rock surfaces of the inner continental shelf. Drowning at times of high interglacial sea levels, and in the Holocene, give this section of coast a deeply embayed appearance. The vertical cliffs of the Central Province are cut into  sandstones of different age (lower Permian to midTriassic). Cliffs up to 100 m high are undergoing active erosion. However, the rate of erosion is slow enough to allow houses to be constructed near cliff edges. Recession

New South Wales

of cliffs truncating old valley sides appears to have been underway since the late Tertiary, although this is difficult to establish. In places, they are cut by vertical-sided clefts formed by the weathering and erosion of basaltic dykes. Geophysical surveys and shallow drilling reveal a veneer of sand over the inner shelf to water depths of about 80 m, with exposure of sandstone outcrops forming what may be an extensive marine abraded surface, over 500 km long and 10–15 km wide. This surface was reworked over a range of sea levels during the Quaternary. Dyke rocks protrude above their surface. At the base of the cliffs are almost horizontal shore platforms. These platforms have been the subject of much observation and speculation, especially as to their genesis under sea levels below, at and above present sea level. Typically, they show the influence of interbedded sandstones and shales, with steps and beaches formed in rocks of slightly greater resistance. Processes of wave quarrying, abrasion, and water-layer levelling have been variously invoked to explain aspects of platform morphology. At a number of locations around Sydney (such as Long Bay and Parsley Bay) remnants of what could have been active shore platforms at times of interglacial high sea levels may be observed. Radiometric dating of iron-rich crusts on the surface of some platforms indicates that parts of these platforms predate the Holocene stillstand, as was noted below by the burial of such features in the Southern Province. The possibility that platform surfaces in Central and Southern

⊡⊡ Fig. 21.1.8 Stanwell Park, New South Wales. (Courtesy Geostudies.)

21.1

provinces have been swept by giant tsunamis has also been suggested by Bryant, Young, and others (Bryant et al. 1996). These authors see tsunamis as a significant process in carving the relief of platforms and moving large rocks. However, such a process has not been observed on this coast, although massive rocks have been moved by storm waves on platforms near Sydney in historic times. This occurred extensively, during the storms of 1974, on a number of platforms where rocks from below mean sea level were detached and rolled across horizontal benches situated at or above the high water mark. Long gently-curving sandy beaches, shaped by refracted southeasterly swell, continue along the central coast of New South Wales. At Stanwell Park (>Fig. 21.1.8) changes on the beach over the past century have been analysed to examine the relationship between erosion and accretion and fluctuations in the beach water table (Bryant 1985). Sydney Harbour (Port Jackson) is a classic example of a drowned river valley or ria system. Ancient dendritic valleys with vertical cliffed slopes have been carved in Triassic rocks (Hawkesbury and Narrabeen groups). Quartzose sandstones predominate with bedrock valley depths, extending 40–60 m below sea level. The history of sedimentation is largely confined to the Holocene marine transgression and ensuing stillstand, as in the Hawkesbury River-Broken Bay estuary. As in Port Hacking and Sydney Harbour to the south, there is a tripartite separation and inter-relationship of three sediment types: flood-tide delta

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New South Wales

⊡⊡ Fig. 21.1.9 Deltas (silt jetties) of Wyong Creek (left) and Ourimbah Creek, built into Tuggerah Lake. (Courtesy Geostudies.)

sands, central basin muds, and fluvial delta sands and muds (Roy et al. 2001). Some lagoons have elongated deltas built by inflowing rivers (>Fig. 21.1.9). Marginal prograded sand barriers are also present, adjoining the flood-tidal delta (as at Woy Woy). There are also coastal lakes, blocked on the seaward side by a sand barrier. For the most part, the barriers are of Holocene age, but Pleistocene dunes complicate barrier morphology in some places. Lake Illawarra, Lake Mac­quarie, and Tuggerah Lake are three of the larger lakes in the province, although several others occur near Sydney (e.g. Narrabeen) and farther south. Botany Bay is somewhat different, with an extensive sand barrier to the south, but a deep and wide entrance between the Hawkesbury Sandstone cliffs to the east permits storm waves and swell to sweep inland across the sandy bay floor and form a beach ridge plain of Holocene age behind Lady Robinson’s Beach. Jervis Bay, towards the south of the province, is a similar embayment, but here the sand barrier occurs on the northern side, and the entrance between the high cliffs to the east, cut into sandstones of Permian age, allows swell waves to penetrate across the width of the bay. Just to the north of Jervis Bay is an excellent example of a mature river estuary, the Shoalhaven. The morphology and stratigraphy of this estuary has much in common with other larger river systems on the New South Wales coast. A mid-Holocene lagoon formed behind a prograded barrier

was filled with fluvial muds and sands from the west, and marine tidal delta and barrier sands from the east.

4.  Northern Province Fold belt terrain, together with rocks of the ClarenceMoreton Basin in the far north into which was extruded the great Mount Warning volcanic caldera of Tertiary age, forms the bedrock framework for the Northern Province. The various rock types have characteristic weathering features and morphology, as on the polygonal basalts at Fingal Point (>Fig. 21.1.10). In contrast to the southern and central provinces, the coastline north of Newcastle extending into southeast Queensland is predominantly sand-rich. Embayments are commonly filled with sediment, so that the space available for marine and fluvial deposits is mostly occupied. Along the open ocean coast, sand bypassing occurs, where in the other provinces the dominant tendency is for embayments to function as ‘closed’ compartments. The overall pattern of sedimentation is for multiple bay-barriers, dune-fields of Pleistocene as well as Holocene age, and extensive fluvial deposits overlying estuarine and pro-delta accumulations. The latter form layered sequences within embayments. There are examples of late Quaternary interglacial/interstadial features at or below sea level (Roy et al. 1994). River terrace-remnants of pre-Holocene age

New South Wales

21.1

⊡⊡ Fig. 21.1.10 Fingal Point, northern New South Wales. (Courtesy Geostudies.)

also occur within the valleys, although their morphostratigraphic histories are not unequivocal. Shore platforms and cliffs, whilst quite striking in some places (such as near Coffs Harbour), have not attracted the same attention in this province as elsewhere in New South Wales. Extensive bay-barrier sequences have been the subject of mapping, drilling, and dating at a number of locations in the Northern Province, notably in Port Stephens, the Myall Lakes, and Tuncurry, with reconnaissance drilling and mapping in the Clarence, Richmond, and Tweed embayments. Detailed coastline evolution studies have been undertaken at Woody Bay near Illuka on the far north coast (Goodwin et al. 2006). Patterns of deposition are repetitive from one embayment to another, providing a level of confidence in any interpretation of the evolution of coastal landforms (Thom et al. 1992). The occurrence of dual (inner and outer) barriers, first reported in New South Wales in the Port Stephens-Myall Lakes area, became the basis for extrapolation of the model of two episodes of accretion, one during the Last Interglacial and the other accompanying the Holocene stillstand. Four sites depict the diversity of bay barrier formation in the Northern Province. The first two are in the Port Stephens-Myall Lakes area and were described in detail by  Thom et al. (1992). Dual barriers occupy the New­ castle  Bight embayment. Reworked dunes formed on an interglacial inner barrier during the Last Glacial Maxi­ mum. Uranium-series dates were obtained on corals from

beneath the inner barrier, the first radiometric age indication of Last Interglacial origin. The second site occurs near Seal Rocks, where a massive transgressive dune sheet, now forested, abuts Myall Lake. Holocene dune sands have migrated landward across a mostly buried barrier of the presumed Last Interglacial age. Radiocarbon dates from buried soils within the dune sheet reveal a late Holocene age of dune migration (less than 2,000 years for the latest phase) (Thom et al. 1992). The site highlights how waves of dune sand can migrate, become stabilised, then be activated again. Tuncurry on the Mid North Coast is the third site, which has a complex of marine sand deposits covering two Pleistocene interglacial events, several interstadials, and the Holocene. A headland-attached shelf sand body lies to the south of the barrier complex, which occurs both onshore and on the inner continental shelf. Holocene ages were determined using radiocarbon dating, and Pleisto­cene dates were based on thermoluminescence measurements. This site clearly demonstrates a pre-Last Interglacial barrier system. From seismic studies through the axis of the embayment, a drowned regressive barrier has been interpreted, based in part on the concordance between the attitude of dipping reflectors and the present shoreface profile. The fourth site is far to the north, in the vicinity of Evans Head. As with other sites, the embayment is filled with sediment, but here the Pleistocene barriers form a  strand plain occupying all the available space. No

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New South Wales

Holocene barriers are present. Sandrock (humate-impregnated sand) outcrops on the beach face as irregular cliffs. Just to the south of this embayment late Holocene infill of small compartments subject to longshore sand transport indicates changes in coastline alignment with changes in wave climate (Goodwin et al. 2006). Studies of estuary and flood plain deposits of the many large rivers in the Northern Province highlighted the diversity of fluvial deltaic landforms (Troedson and Hashimoto 2008). More recent work has concentrated on sub-surface investigations and the relation between estuary deposition and acid sulphate soil formation. Roy et al. (2001) demonstrated a history of estuary infill forming mature river estuaries behind bay barriers.

5.  Conclusion The coast of New South Wales thus possesses a rich array of depositional and erosional landforms. Concerted efforts by a number of investigators have unravelled the history of the continental margin, including coastal evolution from late Mesozoic times to the present. It is a coast that has provoked many questions and various degrees of controversy. Different interpretations of sea level from shore ­platforms extend back to the 1930s, and there are still unresolved issues of detail in the Late Quaternary, especially during the so-called Holocene stillstand. However, first-order patterns are reasonably well established, and it is these patterns that appear to control gross morphological and depositional histories. There are also differences of emphasis on processes and resulting landforms, especially the possible role of tsunamis (Bryant et al. 1996).

References Bird ECF, Dent OF (1967) Shore platforms on the south coast of New South Wales. Australian Geogr, 10:71–80 Bryant EA (1985) Rainfall and beach erosion relationships, Stanwell Park, Australia 1895–1980. Z Geomorphol, Suppl (Bd) 57:51–65 Bryant EA, Young RW, Price DM (1996) Tsunami as a major control on coastal evolution, southeastern Australia. J Coastal Res, 12:831–804 Chapman D, Geary M, Roy PS, Thom BG (1982) Coastal evolution and coastal erosion in New South Wales. Coastal Council of N.S.W., Sydney Goodwin I, Stables M, Olley J (2006) Wave climate, sand budget and shoreline alignment evolution of the Illuka-Woody Bay sand barrier, northern NSW, Australia, since 3000 yr bp. Mar Geol 226:127–144 Langford-Smith T, Thom BG (1969) Coastal morphology of New South Wales. J Geol Soc Australia, 16:572–580 McLean R, Shen J-S (2006) From foreshore to foredune: development over the last 30 years at Moruya Beach, NSW, Australia. J Coastal Res 22:28–36 Roy P, Cowell P, Ferland M, Thom B (1994) Wave-dominated coasts. In: Carter R, Woodroffe C (eds) Coastal evolution. Cambridge University Press, Cambridge, UK, pp 121–186 Roy PS, Thom BG (1981) Late Quaternary marine deposition in New South Wales and southern Queensland – an evolutionary model. J Geol Soc Australia, 28:471–489 Roy PS, Williams RJ, Jones AR, Yassini I, Gibbs PJ, Coates B, West RJ, Scanes PR, Hudson JP, Nichol S (2001) Structure and function of south-east Australian estuaries. Estuar Coastal Shelf Sci, 53:351–384 Short AD, Trembanis AC, Turner IL (2000) Beach oscillation, rotation and the Southern Oscillation, Narrabeen Beach, Australia. Coastal engineering 2000, American Society Civil Engineers, Reston, VA, pp 2439–2452 Thom B (1984) Transgressive and regressive stratigraphies of coastal sand barriers in southeast Australia. Mar Geol 56:137–158 Thom BG, Shepherd M, Ly CK, Roy PS, Bowman GM, Hesp PA (1992) Coastal geomorphology and quaternary geology of the port StephensMyall lakes area. Department of Biogeography and Geomorphology, Australian National University, Monograph No. 6, Canberra Troedson A, Hashimoto R (2008) Coastal Quaternary geology north and south coast of New South Wales. Bulletin 34, Geological Survey of New South Wales, p 94

21.1.1 Lord Howe Island – (New South Wales)



Eric Bird

1.  Introduction Lord Howe Island lies about 800 km north-east of Sydney in the Tasman Sea, between latitudes 31° 30' and 31° 36' E and in longitude 159° 05' E. It is about 10 km long and up to 2.6 km wide, with two high volcanic mountains at the southern end, Mount Gower (875 m) and Mount Lidgbird (777 m), a narrow central isthmus with a broad lagoon and coral reef on the western side (> Fig 21.1.1.1), and a hilly northern part, rising to a north-facing escarpment cliff up to 200 m high (> Fig  21.1.1.2). The coastline is about 25 km long, the southern mountains bordered by precipitous cliffs of volcanic rock, while the central isthmus has some sandy and gravelly beaches and lower cliffs with rock platforms. The surrounding sea is relatively shallow, the sea floor being a broad subhorizontal surface cut by Plio-Pleistocene marine abrasion.

Mean annual rainfall is 1,800 mm, and there have been occasional downpours, as in 1996 when over 420 mm of rain fell in a day, causing landslides and gulleying on steep slopes. Much of the island is wooded, but there are areas of grassland. Mean spring tide range, measured at the Jetty on the west coast, is about 1.5 m, diminishing to 0.8 mat neap tides. Winds, often strong, include S to SE trade winds mixed with westerlies that prevail in winter. Ocean swell arrives from all directions, but a SSW swell dominant on the west coast arrives from south of Tasmania and perhaps through Bass Strait, and the island also receives occasionally strong wave action from the north-east and south-east. The Late Tertiary volcanic rocks include the Roach Island tuff (interbedded tuffs with thin lavas and beds of coarse fragmental material) overlain by many lava flows of the North Ridge basalt (Old Gulch). The Boat Harbour

⊡⊡ Fig. 21.1.1.1 The world’s southernmost coral reef (extending to 31° 34' S) on the west coast of Lord Howe Island, looking south to the volcanic mountains, Mount Lidgbird and Mount Gower. Waves are breaking on the reef edge, and the shallow lagoon is backed by a sandy beach and low cliffs cut in dune calcarenite. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.1.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.1.1

Lord Howe Island – (New South Wales)

⊡⊡ Fig. 21.1.1.2 An escarpment cliff in southward-dipping lava and breccia on the north coast of Lord Howe Island, looking west from Malabar Hill (209 m) past Kim’s Lookout (182 m) to Mount Eliza (147 m). (Courtesy Geostudies.)

marine erosion, forming a broad shore platform, which passes seaward into the coral reef. The present fringing reef is notable for being the southernmost coral reef in the world, on a coast favored by a warm ocean current flowing south from the Coral Sea (Guilcher 1973).

2.  Southern Cliffs

breccia (angular fragments in a massive fine-grained matrix) is overlain by the Mount Lidgbird basalt (lava flows up to 30 m thick, some columnar), which forms the southern mountains (McDougall et al. 1981). Pleistocene dune calcarenite (the Ned’s Beach Cal­ carenite) is extensive in the central part of the island, forming a higher ridge along the eastern side. It consists of a white or yellow-brown sandstone composed of small fragments of coralline algae and lesser amounts of broken coral, shells, and foraminifera, and was formed by erosion of sea floor deposits, including disintegrating coral reefs, by wave action as the sea rose and fell, and the drifting of sand thus derived by wind action. Exposures show laminations inclined in various directions, indicating irregular inland migration of dunes. A ridge of dune calcarenite along the western coast has been cliffed and dissected by

Southward from Little Island on the west coast, a cliff 200– 300 m high, cut in basaltic lava, descends to a basal wooded apron of fallen debris. Little Island is a large tilted block of lava that has fallen to the shore, with a narrow incipient shore platform cut across tilted thinly bedded basalt. There is a storm-piled beach of well-rounded large grey basalt and white coral cobbles and boulders where the southern end of the coral reef comes inshore. To the south, a pale scar on the cliff face indicates the site of a rock fall, with angular blocks on the shore below a talus apron slope that has revegetated. Further along the lava cliff descends to, and below, the sea, with only a narrow grassy strip part way up. To the south, the cliffs increase in height to more than 500 m, and are subject to parallel retreat, maintained by basal marine erosion of the talus apron (known as the Little Slope). Precipitous cliffs of volcanic rock continue round the southern end of the island, past South Head to King Point and outlying steep-sided Gower Island, and round to East Point. The rocky shore is irregular and bouldery, with no platforms. There is said to be a dune calcarenite deposit on volcanic scree here. Offshore, 23 km to the south-east, is Ball’s Pyramid, a jagged residual volcanic stack rising abruptly from the ocean to a height of 551 m. The south-eastern coast has cliffs at the base of steep talus on the Big Slope. Red Point is an elongated promontory of weathered volcanic rock, and the coast north from East Point has several narrow inlets cut out along the strike of the numerous NE-SW dykes that radiate from Mount Lidgbird. Boat Harbour Cove is a north-facing horseshoe bay containing a curved beach of grey basalt and white coral pebbles cobbles heaped in a berm, with an upper storm beach of cobbles and boulders. Pale nearshore areas are sandy with coral boulders, the remains of a disintegrating reef. Live coral reefs exist along parts of the east coast, and the boulders have been thrown up during occasional severe easterly storms. Seaward, the bay frames a view of high grassy Mutton Bird Island and its smaller neighbour, a basalt platform surrounding a residual stack known as Sail Rock, an Old Hat type of shore platform.

Lord Howe Island – (New South Wales)

Rocky Run is the mouth of a deep gully, a cove cut into black outcrops of breccia alternating with basaltic lava, behind which a steep slope ascends to cliffs of lava and basalt columns on Mount Lidgbird. The coast north from Rocky Point has a steep wooded slope down to a black rocky shore, with some structural benches. Offshore, rough white water breaks around a submerged rock. Mutton Bird Point is an almost breached promontory of basalt, with dark zones of underwater platform. There are hillside outcrops of dark breccia, and several landslides that occurred during a major downpour in June 1996 when 420 mm of rain fell in 8 h. The slumped areas have grassy scrub, in contrast with the general Pan­ danus woodland. The volcanic cliffs decline northward to Blinkenthorpe Bay, where Blinky Beach is a sandy surf beach backed by a cliffed dune ridge with grassy and rushy vegetation. An air photograph taken some time after 1974 (when the airstrip was completed) showed bare dunes here. Between Blinky Point and Clear Place Point are cliffs and bluffs of basalt 10–20 m high, with a narrow basal rock ledge, not obviously related to any particular lava flow (similar to that on the north coast). Below Clear Place, there is a natural arch and blowhole, and in the cliff face a large cavity, which seemed to have been in the original volcanic rocks, rather than a weathering feature. There is a  prominent rock ledge 3–4 m above high spring tide level,   probably cut during a Last Interglacial high sea level (Woodroffe et al. 1995), separated by a vertical cliff ⊡⊡ Fig. 21.1.1.3 High tide shore platform cut in basalt, Clear Place Point. (Courtesy Geostudies.)

21.1.1

from the modern high tide subhorizontal shore platform (> Fig 21.1.1.3). On the headland, the basalt shore is irregular but with planed segments, some of them on scoriaceous horizons, and to the north a broad rocky shore platform on breccia bordered by ridges of weathering rock. Some of the breccia is gravelly.

3.  Middle Beach South of Middle Beach, the dune calcarenite forms a steep wooded slope down to a shore with outcrops of black volcanic rock. The beach backed by the dune calcarenite bluff has a ramp of basalt and coral conglomerate, evidently cemented by calcareous seepage. Similar seepage has formed travertine skins across the basalt shore, which have planation corridors and dyke ridges. Beside the southern steps, there is brown basaltic clay in the cliff base, and a gully that was eroded during the 1996 downpour, when soft unconsolidated sand was swept out beneath a breached calcarenite crust and washed down to form a fan on the shore, soon dispersed by wave action. Middle Beach has cliffs, stacks, and shore platforms cut in Pleistocene dune calcarenite. The outcrops include laminated sandstone with bird’s foot imprints, layers with root concretions, and a basal buff fine sand with some palaeosol development, above which is a thick white calcarenite in which are vertical columns (diameter about 10 cm) that had been occupied by trunks of palm trees

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Lord Howe Island – (New South Wales)

rooted in the underlying brown sandy soil, overrun (rather quickly) by an advancing dune. There is karstic weathering of lapies, pitted and honeycombed rock, solution basins and scoured potholes, notches and visors, freshly abraded sloping ramps (and some with green seaweed) down to a flat low tide bench with a network of low algal ridges enclosing shallow pools, making “platforms a vasques.” Besides sand, the beach contains pebbles and cobbles of calcrete and coral, derived from an offshore reef. North of Middle Beach, the coastal scrub is heavily wind-pruned, into a convex curve. The coastal slope has craggy outcrops of grey dune calcarenite. The low tide to subtidal shore platform cut in dune calcarenite extends north from Middle Beach, with segments of emerged shore platform, and a hazardous low level cave overhung by calcrete. Further north is Jim’s Point, a promontory of dune calcarenite showing two sets of dune bedding and a shore platform exposed at low tide. North of Stevens Point, Hell’s Gates are a narrow cove floored with a basalt shore platform diversified by hard dykes and other residuals, fronting a cliff in which bright red rock (baked tuff) is interspersed with dykes. Some of the dykes are resistant and protrude; others, more fractured, have been excavated as ditches or clefts, and the platform is strewn with well-rounded boulders and cobbles of basalt. Searles Point, to the north, is a low promontory of dune calcarenite, cliffed on its southern side and intricately dissected by solution to form rugged lapies

along its northern flank, which is frequently sprayed with sea water as waves break along the shore.

4.  Ned’s Beach North of Searles Point, the dune calcarenite cliffs have lapies that are very hard and spiky, and there are small rounded coves with beaches of calcareous sand and gravel. The initial rate of muricate weathering on a fresh calcarenite surface may be rapid over a few decades, but when it hardens the rate must be much slower. They are fronted by a shore platform, and pass laterally into cliffs exposing 4 m of brown sandy sediment, with many shell fragments and pieces of turtle bone and birds, overlain by at least 18 m of pale grey dune calcarenite. The brown sandy sediment was interpreted as a basal beach by Woodroffe et al. (1995), dated by the uranium series method at 136,000 years bp, supported by thermoluminescence dates in the range 116,000–38,000 years bp, i.e. Last Interglacial, formed when the sea stood 2–4 m above present level. It is possible that it was a lagoonal deposit, landward of a dune calcarenite barrier formed when the sea was higher in Late Pleistocene times. Remains of this enclosing barrier persist on the shore as a dissected karstic ridge of dark dune calcarenite, with another one planed off to seaward. The dune calcarenite contains turtle bones, bird remains, and darker patches that may have been bird sites. The beach or lagoonal deposit is overlain by Late Pleistocene to ⊡⊡ Fig. 21.1.1.4 Arcuate outcrops of submerged beach rock (1 m below low spring tide) off Ned’s Beach, indicating former positions of the beach. (Courtesy Geostudies.)

Lord Howe Island – (New South Wales)

21.1.1

Holocene bedded dune calcarenite, harder and darker in the cliff face but softer and paler, with little obvious stratification, in sectors where the outer weathered surface has been removed. The curving sandy beach is backed by low grassy dunes. There has been progradation during the past few decades, when a former dune cliff dropped to a bouldery shore now covered by beach and dune sand with grasses. The bay is exposed to a north-easterly ocean swell. On the  sea floor are submerged outcrops of beach rock, ­indicating that the beach was formerly much wider here (> Fig 21.1.1.4). Behind the northern part of the of beach, bluffs of dune calcarenite gives way sharply to cliffs cut in irregular tuff with basalt dykes (the boundary is somewhat south of that shown on the geological map) The northern shore is lined with basalt boulders rather than shore platforms, but there is a shore platform and rock residuals in breccia on the north point of Ned’s Beach.

tropical coral lagoon rates) until 4,000 years ago, followed by slower accretion (Woodroffe et al. 2000). Behind it is a sandy Lagoon Beach with low parallel grassy beach ridges, truncated by a cliff up to a metre high on the lagoon shore. Lagoon Beach is interrupted by a sector of sloping sea wall and a narrower beach north of the airport run­ way protrusion. Toward the southern end of the lagoon, the sandy beach is interrupted by bouldery areas below ­former landslides. There is sporadic coral growth some way out in the lagoon, but the true reef garden is to seaward, and the reef ends in a 100 m seaward decline to a plunging edge. The fringing coral reef has developed on and seaward from a planed-off dune calcarenite platform. There are large holes or depressions in the reef, with ledges and seagrasses as well as various fish. Acanthaster is present but sparse – whereas on Elizabeth and Middleton Reefs it has been devastating, and shows cyclic revival. There is some coral

5.  North Coast

⊡⊡ Fig. 21.1.1.5 Escarpment cliff at Malabar Hill. (Courtesy Geostudies.)

The north coast is an escarpment cliff rising to 209 m at Malabar Hill and Kim’s Lookout (182 m) (> Fig 21.1.1.5). Below is Soldiers Cap, an Old Hat island with a surrounding high tide shore platform cut in tuff, and beyond that the Admiralty Islands, also cut in tuff, with a natural tunnel on Roach Island. The high basalt cliff on the north coast plunges to a narrow basal ledge, then deep water inshore. where the northern coast is an escarpment cliff on southwarddipping lava. To the west Old Gulch is a valley-mouth inlet between basalt cliffs and undulating high tide ledges cut in breccia and across dykes, backed by a curved storm-piled beach of basalt and coral cobbles and boulders with a few tumbled blocks of dune calcarenite. Beyond it, the escarpment cliffs rise again to Mount Eliza (147 m).

6.  West Coast Lagoon The Lord Howe Island coral reef is about 6 km long and extends to just below 31° 34' S. Surf generated by a southwesterly swell breaks across the coral reef rampart, which has splays of dark gravel, and the reef descend seaward into deep water by way of a slope with large coral patches amid sandy areas. The reef crest, formed between 5,000 and 6,000 years ago, is backed by a broad lagoon, shallow at low tide, with a rocky floor and sandy and muddy patches, produced by relatively rapid deposition (comparable with

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21.1.1

Lord Howe Island – (New South Wales)

bleaching, but white tips of coral are also caused by snail attacks. The corals are replenished by polyps carried in the East Australian Current. There is an abundance of soft algae, particularly Caulerpa sedoides, indicating a transition from tropical coral reefs to temperate algal communities (Guilcher 1973). Within the lagoon is Blackburn (Rabbit) Island, a grassy basalt island about 10 m high, with dykes exposed in low cliffs and some gravelly beaches. At the northern end of the lagoon is North Bay containing North Beach, which curves round to the northern end of the coral reef, and the inner platform on dune calcarenite has aggraded into an outer coral reef. At the western end is a boulder spit, with bleached white coral boulders above high tide and brown below, running out to the northern end of the reef rampart (> Fig 21.1.1.6). At the eastern end is a cove with a cobble beach and some outcrops of dune calcarenite, which on a dissected headland showed spiky lapies, visors, and notches, with ramps below scoured by cobbles. The coralline sand coarsens to grit as the beach curves westward, backed by a dune cliff and grassy ridges. In the low tide lagoon shallows are subdued shoals with sparse seagrasses.

7.  Old Settlement Beach East of the wooded promontory of Dawsons Point is Hunter Bay, backed by Old Settlement Beach, of coarse

coralline sand with some coral gravel. There is a single mangrove on the western shore. Sylphs Hole is prominent in the lagoon offshore. The creek at the eastern end of Old Settlement Beach is bordered by several mangroves (Avicennia) beside a bridge, and an area of salt marsh. On the eastern shore is a group of planed-off dune calcarenite blocks, formed by dissection of an emerged mid-Holocene shore platform about 1.5 m higher than the present, cut when low tide was just above present high tide line. A flat-topped mushroom rock standing above the platform shows the past and present levels of planation, and on the coastal slope there are karstic knobs of dune calcarenite in woodland. At the Jetty (completed in 1983) north of Flagstaff Hill, there is an instrumental tide gauge. A sector of low cliff in soft dune calcarenite has been armoured with a sloping apron of basalt boulders. Deepening of the nearshore area near the Jetty by ship scour exposed a clay horizon 2 m below sea level with the remains of mangroves, dated about 6,100 years bp.

8.  Signal Point Signal Point at Flagstaff Hill is a headland of dune calcarenite with rugged stacks and mushroom rocks on a low tide shore platform passing under the lagoon. There are dissected remnants of a mid-Holocene emerged shore platform. Below Signal Point (Flagstaff Hill) are several

⊡⊡ Fig. 21.1.1.6 Coral gravel spit, North Beach, looking toward Mount Gower and Mount Lidgbird. (Courtesy Geostudies.)

Lord Howe Island – (New South Wales)

21.1.1

⊡⊡ Fig. 21.1.1.7 Pleistocene dune calcarenite at Callaghan’s Beach on the west coast. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.1.8 Shingle beach toward the southern end of the west coast lagoon, looking south to the high cliffs in volcanic rock rising to Mount Gower. (Courtesy Geostudies.)

irregular dune calcarenite stacks notched at mid-tide level and the planed-off dune calcarenite is extensive at low tide in the lagoon.

9.  Cobbys Corner to Johnsons Beach The lagoon beach extends south to a series of low dune calcarenite promontories with sandy coves. Callaghan’s

Rock is a cliff cut back into a former dune calcarenite ridge, and to the south the sandy beach has outcrops of black basalt but the cliffs are in the overlying dune calcarenite (> Fig 21.1.1.7). At the northern end of sandy Johnson’s Beach is a dune cliff with outcrops of northward-dipping calcarenite. This was part of a broad dune calcarenite ridge on the eastern side of the present lagoon, running north–south and now much dissected, with cliffy headlands, stacks, and shore

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Lord Howe Island – (New South Wales)

platforms. There is evidence of an earlier phase of planation just above present high tide level. A small creek with some mangroves (mainly Avicennia) opens to the shore at Johnson’s Beach, and has inwashed calcareous sand and reefs of planed-off knobbly dune calcarenite with many stacks and cliffy headlands. The beach becomes bouldery as the volcanic rocks come to the coast, and the reef and lagoon narrow southward to Little Island (> Fig 21.1.1.8).

References Guilcher A (1973) Lord Howe, l’île a recifs coralliens la plus meridonale du monde. Bull Assoc Géogr 404–405 and 427–437 McDougall I, Embleton BJ, Stone DB (1981) Origin and evolution of Lord Howe Island, south-west Pacific Ocean. J Geol Soc Aust 28:155–176 Woodroffe CD, Murray-Wallace CV, Bryant EA, Brooke B, Heignis H, Price DM (2000) Late Quaternary sea-level high stands in the Tasman Sea: evidence from Lord Howe Island. Mar Geol 125:61–72

21.1.2 Norfolk Island

Eric Bird

1.  Introduction Norfolk Island (29° 02' S, 167° 57' E) lies 1,760 km ENE of Sydney, 900 km ENE of Lord Howe Island, and 675 km south of New Caledonia. The main island (35 sq. km) is about 8 km E-W and 6 km N-S, with a coastline 32 km long, and a highest point (Mount Bates) 318 m above sea level. There are two smaller islands to the south. Nepean Island (4 ha) rising to 32 m, with a 1.2 km coastline and Phillip Island (1.4 sq km) rising to 280 m, with a 6 km coastline (>Fig.  21.1.2.1). The climate is subtropical, with a mean monthly maximum (February) of 22°C, a mean monthly minimum (August) of 16°C, and an average annual rainfall of 1,326 mm. The prevailing winds are easterly (NE–SE) in summer and westerly (S–W) in winter, when gales occur. Waves generated by these winds are accompanied by ocean swells mainly from the SW and SE, while tropical cyclones to the north generate waves from that direction (Jurd 1989). Mean spring tide range is about 1.3 m. Norfolk Island and Phillip Island are of volcanic origin (Jones and McDougall 1973; Abell and Falkland 1991). They rise from the mid-Tasman Norfolk Island Ridge, which runs from New Caledonia southward to Auckland, New Zealand. The two islands consist of lavas and tuffs produced by eruptions in Pliocene times, between 3.3 and 2.5 million years ago, and have subsequently been reduced in size by marine erosion. The volcanic formations include the Ball Bay Basalt and Duncombe Bay Basalt, overlain by the Cascade Basalt and the Steels Point Basalt. The lava flows are up to 30 m thick, and often show columnar jointing. These are separated and overlain by layers of tuff and weathered basalt (clay with boulders) formed on intervening topographic surfaces. The tuffs are up to 15 m thick, horizontally or current bedded, and vertical in cliff sections where they are overlain by basalt. The volcanic deposits are generally flat or gently dipping, indicating little tectonic deformation since they were deposited. The main vents were at Mount Bates and Mount Pitt (316 m) on Norfolk Island and Mount Jacky Jacky (280 m) on Phillip Island. Away from the higher ground of Mount Bates and Mount Pitt, Norfolk Island is dominated by a plateau 100–120 m above sea level, on which the

v­ olcanic formations have been deeply weathered, in places to depths of 45 m, and there are reddish brown krasnozem soils with scattered spheroidal corestones. These occur on the surface locally, and are an already-rounded component of some boulder beaches. The plateau is dissected by several river valleys, but there are few perennial streams, and apart from Watermill Creek, which descends to sea level at Kingston, they reach the coast in hanging valleys with waterfalls, indicating that cliff recession has been more rapid than valley incision. The absence of valleys incised below present sea level means that there are no inlets resulting from Late Quaternary marine submergence and hence no natural harbours. The coast of Norfolk Island is generally steep. On weathered basalt and tuff, the steep (up to 50°) slopes are vegetated and subject to landslides, while the harder solid basalt has been cut back in vertical cliffs. Slope-over-wall profiles are common where a vegetated slope in weathered basalt and tuff descends to a vertical cliff in basalt (>Fig. 21.1.2.2), the morphology of which is much influenced by vertical joint planes. Shore platforms are generally structural, formed where the surfaces of resistant lava beds have been exposed on the shore, but locally subhorizontal weathered platforms have formed on tuff and weathered basalt. Along the north coast, the Pliocene lavas were deposited in a contemporary sea, and show a quench zone (hyaloclastic deposits) extending up to 15 m above the present shore, showing pillow lavas and hackly columns formed by rapid chilling by sea water. On these shore platforms are irregular surfaces, and shattered rocks have disintegrated into boulders. On the south coast of Norfolk Island at Kingston, the cliffs pass inland as grassy bluffs, which were coastal cliffs during the Pleistocene. Slopes of this kind doubtless developed around the island during low sea level stages, and were rejuvenated as cliffs as the sea rose to its present level in the Late Quaternary marine transgression. The Kingston bluffs overlook a lowland dominated by Pleistocene dune calcarenite, laminated calcareous sandstones of aeolian origin. The dune calcarenite has been dated at about 22,000 years bp, and formed during a Pleistocene low sea level

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.1.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 21.1.2.1 Norfolk Island and Phillip Island. (Courtesy Geostudies.)

stage when calcareous marine sediment, including shell and coral fragments, were blown up from an emerged sea floor by southerly winds. Dune calcarenite outcrops in low rugged cliffs at Point Hunter, with notches and visors and karstic weathering features fronted by shore platforms exposed at low tide, which have been produced by solution and bioerosion

processes as well as wave abrasion. The shore platforms continue as low tide reefs west from Point Hunter to the Kingston Jetty, and colonization of these by coral and algal colonies has produced an incipient fringing reef. Nepean Island, offshore to the south, also consists of a cliffed ridge of dune calcarenite bordered by low tide shore platforms.

Norfolk Island

⊡⊡ Fig. 21.1.2.2 Slope-over-wall coast at Anson Point. (Courtesy Geostudies.)

  21.1.2

as on the north shores of Cemetery Bay, in Ball Bay (where there are also pale coralline cobbles), at Cascade and south of Anson Bay (>Fig. 21.1.2.3), where columnar basalt dips down to the shore Boulders fronting some basalt cliffs, while others have shore platform segments or plunge directly into the sea.

2.  The Island Coastline

Behind the dune calcarenite fringe, Watermill Creek flows into swampy areas, which are the remains of a lagoon that existed behind a dune-capped barrier. Deposits of black carbonaceous clay in the former lagoon behind Slaugh­­ ter  Bay and Cemetery Bay have yielded Norfolk Island pine logs with radiocarbon dates of about 4,400 years. Sandy beaches are rare on Norfolk Island because the weathered basalts and tuffs have not yielded sandy sediment: there are no black or grey beaches of pyroclastic origin like those of Hawaii or Indonesia. Beaches of calcareous shelly and coralline sand swept in from the sea floor and generated from the weathering and erosion of dune calcarenite are found in Slaughter Bay (where there are layers of beach rock), in Emily Bay, east of Point Hunter and in Cemetery Bay. They are backed by low dunes in Emily Bay and Cemetery Bay. There are also small beaches of calcareous sand in Creswell Bay to the west, and Anson Bay on the northwest coast. Elsewhere, there are beaches of ­well-rounded blue-grey basalt cobbles and pebbles locally,

There is public access to about one-sixth of the coastline, but much is private land. There is a great need for a coastal footpath to make this magnificent coastline accessible. Kingston Jetty, on the south coast, is a stone structure built beside a gap in the reef of dune calcarenite, which extends eastward to Point Hunter. There is often a strong SW swell here, breaking alongside and over the jetty. The steep and cliffy basalt coast that dominates Norfolk Island here passes into a grassy bluff behind the Kingston lowland, which has an impressive array of historical buildings dating from the nineteenth century penal settlement. Watermill Creek flows out of its valley, across a grassy area maintained by cattle grazing and through a sedgy wetland and a short canal to the sea at Emily Bay. East of Kingston Jetty, a sea wall is fronted by an abrasion ramp of Pleistocene dune calcarenite, then a low tide shore platform cut in dune calcarenite. There is a sandy beach in Slaughter Bay, where at low tide multiple layers of beach rock are exposed. These formed during a period of beach progradation, and have been exposed as the result of beach depletion. Salt House headland has the remains of lime kilns beside a quarry in the dune calcarenite. Near the Salt House, an emerged dune calcarenite platform contains basins that were excavated to trap and concentrate sea water to brine before salt extraction. The emerged platform ends in a 1–2 m cliff dropping to the present low tide shore platform. Behind the headland is an archaeological site with artefacts that indicate a phase of Polynesian occupation around 1200–1600 ad. The dune calcarenite reef exposed at low tide offshore has scattered corals and some coral colonies, particularly in the lagoon off Emily Bay, where tourists are taken out a glass-bottomed boat to see them. These are not true coral reefs, but coral gardens that may eventually develop into coral reefs. Emily Bay has a curving sandy beach backed by low dunes and pinewoods. The dunes have been cut back to a metre high cliff, protected by a pine log fence at the eastern end. This erosion may have been the result of the cutting of a channel through the dune calcarenite reef off Emily Bay, allowing larger waves to cause backshore dune cliffing.

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⊡⊡ Fig. 21.1.2.3 Columnar basalt in Anson Bay dips down to the shore, where angular boulders have been derived from it. (Courtesy Geostudies.)

The Kingston Lowland runs east through the golf course, bordered by a ridge of dune calcarenite east from Point Hunter. On the seaward side, this has been eroded into rugged cliffs, with notches and visors, karstic pinnacles, and slabs of dark dune calcarenite patchily broken to reveal yellow sandstone. The dune bedding is generally northward, indicating emplacement by southerly winds. The cliffs are fronted by wide shore platforms on dune calcarenite exposed at low tide. These platforms were formed largely by solution processes, including rain and sea spray, which remove the soluble rock down to the level of permanent sea saturation. There is a sharp drop at the seaward edge. In Cemetery Bay, low dunes back a beach of calcareous sand over an emerged dune calcarenite ledge, with a step down to the present low tide shore platform. The calcarenite has been dissected, and there is some calcareous gravel at the southern end. The grassy escarpment to the north descends to an apron of downwashed sediment over an emerged Pleis­ tocene beach. To the east, the escarpment passes into steep bluffs north of Cemetery Bay and rocky cliffs cut in Steels Point Basalt, running east toward Collins Head. There is a beach of cobbles and boulders along the shore and coral reef segments out in Cemetery Bay. Ball Bay is a rounded bay that may have formed as a maar (a crater produced by a volcanic explosion), which became truncated by marine erosion, but the rim of tuff that usually accompanies a maar is not obvious. It is bordered by steep slopes on weathered basalt on which there have been landslides, and by cliffs of solid basalt on Collins Head and

Point Blackbourne, at the bay mouth. On either side, there are truncated spurs of weathered basalt with landslides between deeply incised hanging valleys (>Fig. 21.1.2.4). It is possible that Ball Bay originated as a deep valley in weathered basalt, which was excavated by marine erosion, and owes its rounded form to the refraction of swell arriving from the southeast between ledges and stacks of solid basalt on the headlands. A landslide on the southwestern side is intermittently active, but the slumped material has been rapidly removed by wave scour to maintain the curved coastline. The bordering shores are lined with dark basalt boulder beaches, passing into cobbles at the head of the bay (where wave energy is highest and attrition most rapid), mixed with fragments of pale coral but no sand. The east coast of Norfolk Island, north from Point Blackbourne, has steep slopes on weathered basalt descending to basal cliffs of solid basalt and occasional shore ledges, cut out along joints in the hard lava. It is ­dissected by several short, steep valleys into dells and knolls, and  there have been repeated landslides. To the north, a rounded coastal slope overlooks Steels Point. A steep upper slope on weathered basalt declines to a wall of basaltic lava and a basal outcrop of layered tuff, fronted by a boulder beach (>Fig. 21.1.2.5). Shore platforms have been cut in the tuff (smooth but slightly undulating) and there are ledges and slabs of hard basalt. The north coast of Norfolk Island, west from Steels Point, has steep vegetated upper slopes of weathered basalt with corestones over softer grey tuff then vertical cliffs of black massive lava, with Top Hat an outlying stack. The upper slope recedes behind a sector of grassy bench above

Norfolk Island

  21.1.2

⊡⊡ Fig. 21.1.2.4 The northeast coast of Ball Bay showing truncated spurs and hanging valleys. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.2.5 Steels Point, showing an upper slope in ­weathered basalt, a cliff of unweathered basalt and a basal tuff. (Courtesy Geostudies.)

the basalt cliff, and then declines to the valley of Stockyard Creek, a stream which descends three waterfalls, the last plunging to the shore. To the west of Stockyard Creek, the basalt cliffs have assorted fallen boulders on the shore. A quarried valley ends in gravel and boulders spilling on to the shore near the Cascade Jetty, behind which the slope has been cut back to an artificial cliff, which exposes a sequence of lava flows and intervening weathered basaltic clays (>Fig. 21.1.2.6). A steep winding valley descends to a bay with a beach of grey well-rounded cobbles and

boulders below the site of an old whaling station. Beyond a steep bluff is the larger valley of Cascade Creek, descending over the Cockpit Waterfall to rapids that end in a plunge over knobbly fractured and compressed black lava to the rocky shore. To the west, the coast of Cascade Bay has hanging valleys between steep grassy slopes interrupted by craggy lava flows, with a massive basalt showing columnar jointing along the coast toward Bird Rock. The basalt columns are narrow when compared with the larger structures on thick

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⊡⊡ Fig. 21.1.2.6 Section at Cascade Jetty showing basalt lava flows separated by zones of weathered basaltic clay. (Courtesy Geostudies.)

masses of slowly cooling lava of the kind that formed in Giant’s Causeway in Northern Ireland. The upper slopes rise to a broad bench, backed by higher ground rising to Mount Bates in the forested National Park. Below the cliffs, there are grassy basal aprons of bouldery talus, and the sea floor is littered with large blocks and boulders of basalt, stable enough to have gathered algae, corals and seaweed. Bird Rock is one of the several islands off the north coast, formed by the breaching of headlands and the collapse of natural arches. A little to the east is a long rocky promontory in columnar black basalt, penetrated by a large archway, capped by yellow tuff then more lava flows rising to an upper grey tuff. Nearby is a tall stack crowned with scrub. Elephant Rock also has a natural arch in deformed columnar basalt (>Fig. 21.1.2.7 ), and beyond it Bird Rock is an island consisting partly of columnar basalt, faulted against almost horizontal tuff, with an outer shore platform cut in basalt to the west and tuff to the east. The steep forested coast of the National Park continues westward, basal cliffs of basalt facing out to Cathedral Rock, where the columnar basalt has been pierced by a tall tunnel through which waves wash, Green Pool Stone, which has a shallow depression on a shore platform, and the Moo-oo Stone (Moo-oo is a rush, Cyperus lucidus), a high stack with an outer shore platform. This stands off a grassy spur, on which stands the Captain Cook Monument, recording his arrival here in 1774, and a well-built lookout platform. To the west, the cliffs behind Duncombe Bay rise to a coastal slope on which wind-pruned bent trees indicate easterly gales. There have been landslides on Point Howe, and Point Vincent (>Fig. 21.1.2.8) ends in an

irregular shore platform below a cliff cut in subrounded scoriaceous basalt pillow lavas in a tuffaceous matrix, grading up to hackly jointed lava, a breccia due to the fragmentation, grinding, and quenching of the lava wen it flowed into the Pliocene sea. The west coast of Norfolk Island has bold cliffs in basalt, locally rising to very steep vegetated slopes, as to the north of Anson Point (>Fig. 21.1.2.2). In Anson Bay, a southwest swell rolls in to a sandy beach. The cliff section showed a massive basalt with slightly deformed columns dipping up to 30° south (>Fig. 21.1.2.3), with yellow tuff below and grey tuff above and a capping of red-brown krasnozem. Two fault zones show slumping, and the columnar basalt is downfaulted to sea level at the northern end of the beach. The surf-piled steep beach is backed by a grassy sand terrace, with a metre high cliff cut in the southern part. Although the beach is of similar colour to the yellow tuff, it is of calcareous marine sand with shelly fragments, whereas the tuff is a compact siltstone (lithified by yellow palagonite) weathering to silt and clay sediment too fine to be retained on the surf beach. To the south at Puppy’s Point cliffs of weathered lava descend to a rocky shore with a beach of black cobbles and pebbles. South of Puppy’s Point, a steep upper slope in weathered Steels Point Basalt and buff tuff descends to a  vertical cliff of Cascade Basalt (basaltic lava flows) ­extending to Headstone. In the bay south of Headstone, the vertical cliff of basaltic lava forms a diminishing ­proportion of the coastal profile, the black lava declining to the ledge on Rocky Point where the overlying tuff ­sustains mutton bird burrows.

Norfolk Island

  21.1.2

⊡⊡ Fig. 21.1.2.7 Elephant Rock, showing a natural arch incontorted columnar basalt. (Courtesy Geostudies.)

⊡⊡ Fig. 21.1.2.8 Irregular shore platform and lower cliff on brecciated basalt on Point Vincent. (Courtesy Geostudies.)

Rocky Point Reserve has a mutton bird rookery under scrub on tuff above a basalt cliff and ledges. There are sectors of abrasion ramp cut by storm waves and a circular pool, possibly a lava blister excavated by the sea. On the knobbly weathering basalt platform, boulders are dislodged by marine erosion. To the east steep, bluffs extend round to Point Ross. Crystal Pool is a gap in the shore platforms that have developed where marine erosion has cut along softer weathered zones between lava flows below the western slopes of Point Ross. Weathered basalt outcrops along the cliffs on either side, and to the east Rocky Point Creek valley descends over small waterfalls to Bumbora Reserve and Cresswell Bay. Basalt cliffs and ledges form the bordering headlands, and the valley-mouth cove has been excavated in weathered basaltic clay. There is a beach of calcareous sand behind large black boulders in the sea, the beach containing shelly and coral fragments and patches of exposed beach rock. There is an abrasion ramp in weathered basalt beside the stream mouth. From Cresswell Bay, steep bluffs and basal basalt cliffs with caves, ledges, and stacks run eastward to the gullied slopes behind Kingston Jetty.

3.  Nepean Island Offshore, Nepean Island consists of dune calcarenite, similar to that on Point Hunter and representing the remains of a formerly more extensive foreland. There are basalt

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inclusions in the dune calcarenite, which may overlie a basalt reef. Rugged, dissected cliffs and low tide shore platforms fringe the island, which has a sandy beach in the western bay.

4.  Phillip Island Six kilometres to the south of Kingston, Phillip Island consists of undulating and precipitous slopes of weathered and unweathered basalt and associated tuff, bordered by cliffs. Its formerly rich vegetation cover was destroyed by pigs, goats, and rabbits introduced in the 1790s. Pigs and goats were eliminated by the 1900s, but rabbits persisted, preventing regeneration of the plant cover, until they were eradicated in the 1980s. There has since been some recovery of vegetation, but the island is still dominated by bare gullied slopes in basalt and tuff, and there has been erosion by wind action. On the west coast of Phillip Island, the lava cliffs are up to 90 m high, and in the south hard lava caps tuff in large cirque-like cliffs up to 250 m high below Red Knoll.

A steep-sided narrow isthmus with dykes and sills in tuff, The Razorback, connects the southern peninsula, Jacky Jacky Ridge, rising to 280 m and bordered by steep plunging cliffs. This was the main volcanic vent, the source of lava flows up to 15 m thick that dipped at about 10° northward. On the east coast, cliffs up to 60 m high back narrow bays and promontories such as Juvenile Point, and there are shore benches where waves have removed weathered basaltic clay or tuff to expose the upper surfaces of lava flows. The beaches are of pebbles, cobbles, and boulders. There are several small islands, one of which, Sail Rock, is a volcanic neck.

References Abell RS, Falkland AC (1991) The hydrogeology of Norfolk Island, South Pacific Ocean. Bureau of Mineral Resources, Geol Geophys Bull 234 Jones JG, McDougall I (1973) Geological history of Norfolk and Phillip Islands, Southwest Pacific Ocean. J Geol Soc Aust 20:239–254 Jurd G (ed) (1989) Norfolk Island environment book. Australian National Parks and Wildlife Service. Canberra

21.2 Queensland

David Hopley · Scott Smithers

1.  Introduction The Queensland coastline is 9,880 km long, and extends across 19° of latitude from 28° 10' S at the New South Wales border to 9° 15' S at Torres Strait. A marked reduction in wind and wave energy conditions takes place northward. Significant wave height, exceeded 50% of the time, is 1.15 m at the Gold Coast, 0.8 m near Rockhampton, and 0.5 m at Cairns. Similarly, the occurrence of gale force winds declines from 0.7% at Brisbane to 0.1% at Cooktown. Swell waves dominate the wave spectra as far north as the Tropic of Capricorn, but are progressively excluded to the north by the protecting influence of the Great Barrier Reef, and are totally excluded from the Gulf of Carpentaria. However, the entire Queensland coastline is prone to tropical cyclones during summer months; on average, three cyclones per year occur both in the Gulf of Carpentaria and off the eastern coast. Any part of the coast can experience storms, with central pressures of 960 mb and wind speeds in excess of 150 km/h, but the

coastal zone most prone to intense storms stretches from Princess Charlotte Bay to Fraser Island. Cyclonic seas may raise wave heights to 8 m or more, thus producing occasional high energy events in direct contrast to prevailing conditions. Recent investigations of coastal ridges of detrital coral and shell, located along the northeast Queensland coast and also in the Gulf of Carpentaria, suggest that high intensity cyclones (category 3 or above) occur along this coast, every two to three centuries on average, and have done so throughout the Holocene (>Fig. 21.2.1). Storm surges are also associated with tropical cyclones, and are superimposed on the semidiurnal tidal regime to raise water levels by up to 6 m. Tidal ranges reach a maximum of more than 9 m in Broad Sound on the central coast, but decrease rapidly seaward (ranges not exceeding 6 m on the Great Barrier Reef), southward (a 2-m range being characteristic of southern Queensland), and to the north (a 3-m range being experienced from Townsville to Torres Strait). A 2–3 m range is characteristic of most of the Gulf of Carpentaria, but this rises to 4 m in the

⊡⊡ Fig. 21.2.1 Sequences of coral shingle ridges like these on Curacoa Island, Palm Group, have provided evidence of regular high intensity cyclone occurrence throughout the Holocene.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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southeast, near Karumba. The tides of the Gulf and of the east coast are frequently not in phase, producing complex and unique tidal conditions in Torres Strait, where strong tidal currents in excess of 3 m/s can be experienced. The eastern coast is overlooked by the escarpment of the Eastern Highlands, which rises in places to over 1,000 m. The north-northwest to south-southeast trend of the coastline is determined by the structure of the Tasman geosyncline and its associated basins, which form the axis of the highlands. Cenozoic uplift and a complex history, involving deep weathering and regolith removal, ensure a massive yield of sand- and silt-sized sediment to the east coast. In contrast, the Gulf of Carpentaria represents a more recently downwarped area, the surrounding Mezosoic and Cenozoic sediment forming low-lying, sandy plains with extensive laterite development. Although the region is relatively stable tectonically, Holocene earth movements, possibly of a hydro-isostatic nature, have occurred around the Queensland coast, affecting the Holocene relative sea level pattern and influencing many coastal landforms (>Fig.  21.2.2). Modern relative sea level appears to have been achieved ubiquitously 6,000 years ago, with ages as great as 6,500 years being associated with a modern sea level in some areas. Variations in the maximum height of the sea during the Holocene are indications of recent movements. Some areas of the east coast show no level higher than the present, but large areas retain evidence of 1.5 m or higher. Indication of the recent warping seen is in the maximum level found in the Gulf, from 1.5 m in the north to 2.4 m near Karumba. The most recent review of the climatology, oceanography, and physical features of the coast of eastern Queensland and the Great Barrier Reef is found in Hopley et al. (2007).

2.  Coastal Regions The combination of structural variation, recent history, and variation in energy levels produces a six-fold regional division of the Queensland coastline.

2.1. New South Wales Border to Hervey Bay From the New South Wales border to Hervey Bay, the swell-dominated coast of southern Queensland has massive sand barriers formed by the northward drift of sand from the New South Wales coast. The drift continues as far north as Fraser Island, where it is directed over the continental shelf via Breaksea Spit and into deep water down several submarine canyons. The barriers are of both

Holocene and Pleistocene ages (Ward 2006), but Holocene dunes and ridges have been removed in places, and erosion has exposed extensive areas of Pleistocene humic hard pan on the beach (>Fig. 21.2.3). Australia’s major tourist resort, the Gold Coast, is located in the southernmost section of this barrier coast (>Fig. 21.2.4), where coastal erosion is a serious management issue. North of the Gold Coast a series of massive sand barrier islands have developed, including North Stradbroke, South Stradbroke, and the Moreton Islands, which enclose Moreton Bay. Extensive estuarine flats border Moreton Bay, with beach ridges prevalent in exposed sectors, and mangroves and salt marsh dominating the more sheltered areas. At the northern end of this coastal sector is Fraser Island, which, at more than 122 km in length, is claimed to be the largest sand island in the world. Dunes on Fraser Island rise up to 240 m above sea level, and contain some 40 perched lakes that lie over these dunes. On North and South Stradbroke islands and on Moreton Island dunes rise to 300 m (>Fig.  21.2.5). Most of these large sand masses are Pleistocene in age, and have been weathered into a range of coloured sands.

2.2.  Hervey Bay to Hinchinbrook Island Between Hervey Bay and Hinchinbrook Island a massive sand-size fluvial sediment yield to the coast and a concentrated summer wet season characterise the wide coastal plains of central Queensland. Much of the coast is sandy, with a series of beach ridges trailing northward from the mouth of almost every stream. The coastline is typically made up of up to ten or more ridges up to 500 m wide, but close to major rivers the sequence may widen to over 5 km (>Fig. 21.2.6). Most of the ridges are Holocene in age, but a fragmentary Pleistocene series is found. The only major dune accumulations on this coast occur on Whitsunday Island, Curtis Island, and north of Rockhampton. The largest rivers have built up deltas or large estuarine sediment infills. The Fitzroy River enters the sea in the lee of Curtis Island, and most of the lower estuary consists of bare salt pan (>Fig.  21.2.7) and meandering mangrovelined creeks, a morphology repeated on a smaller scale in most sheltered areas of the coast. The Burdekin River has built Australia’s largest cuspate delta (Hopley 1970) (>Fig.  21.2.8). Under the southeasterly influence, the higher-energy eastern side of the delta has a series of spits  extending from the various distributaries (Pringle 2000). The northernmost, Cape Bowling Green, is more than 18 km long, making it Australia’s longest sand spit (>Fig. 21.2.9).

Queensland

⊡⊡ Fig. 21.2.2 Dates for reef flats on the Queensland coast and Great Barrier Reef.

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⊡⊡ Fig. 21.2.3 Dark humic sandstone (coffee rock) exposed in a section cut in dune sand at the mouth of Currimundi Creek, Sunshine Coast. The humic sandstone is overlain by grey older dune sand on the left, and yellow newer dune sand on the right.

⊡⊡ Fig. 21.2.4 Gold Coast, southern Queensland, just after cyclones Wanda and Zoe in 1974, which removed much of the sand from the beach.

The highest tidal range on the Australian east coast affects the coastline near Mackay, south to Shoalwater Bay, with maximum tidal ranges exceeding 9 m experienced in Broad Sound. Coastal features in this area reflect this high tide range, with extensive mudflats, funnel shaped estuaries, and salt flats having developed on parts of the coast sheltered from the dominant southeasterly winds.

Although high wave energy and lower tidal ranges are normally associated with barrier development, more than 30 barrier systems have been identified along this macrotidal coastline of central Queensland. Only in the south, for example on Curtis Island, do the headlands show any form of shore platform development. Elsewhere, their slopes are essentially weathering forms

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⊡⊡ Fig. 21.2.5 Active dunes on Moreton Island.

⊡⊡ Fig. 21.2.6 Holocene barrier built up from sand from the Burdekin River and consisting of almost 150 ridges, south of Cape Cleveland, near Townsville.

stripped of their regolith. Boulder beaches or boulder spits are found in some bayhead locations (>Fig.  21.2.10). Islands are numerous on this coast, and are drowned extensions of the coastal ranges, continuing the structural trends of the mainland seaward. All but the Northumber­ land Group have extensive fringing reefs. Bay heads are usually occupied by beach ridges. Lee-side sand spits, containing beach rock and older boulder beaches, are very common (Hopley 1971).

2.3.  Hinchinbrook Island to Cooktown This part of the Queensland coast coincides with the highest rainfall region in Australia. Annual totals exceed 2,000 mm throughout the area and reach a maximum of over 4,000 mm near Innisfail. It is predominantly a steep coast, but tropical rainforest extends down to the high tide level (Bird and Hopley 1969). The intertidal zone is rocky, with massive boulders strewn over the foreshore. Close to the

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⊡⊡ Fig. 21.2.7 Bare salt pan (sabkha), northern Curtis Island.

⊡⊡ Fig. 21.2.8 Eastern coastline of the Burdekin Delta showing strong south to north movement of sand from distributaries.

mouths of the major rivers, beach ridges have prograded the coastline by up to 1 km, but there has been erosion of  sectors deprived of a sand supply, as on sectors of the  Barron River delta. Near Kurrimine, a multiple barrier system up to 12 km wide has been formed. The largest  dunes, rising to 60 m, are found on the northern end  of  Hinchinbrook Island. The most extensive mangrove ­systems are found in the Hinchinbrook Channel

(>Fig. 21.2.11) and in Missionary Bay on Hinchinbrook Island, where 31 species have been identified. The rocky shores of the offshore islands are similar to those of the mainland. Many have small, vegetated sand spits on their northern or lee sides. Most have wide fringing reefs, although the largest island, Hinchinbrook, has no reefs, probably because of the large fresh water runoff from the island. Fringing reefs are rare on the mainland,

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⊡⊡ Fig. 21.2.9 Cape Bowling Green, more than 18 km long, is Australia’s largest sand spit.

⊡⊡ Fig. 21.2.10 Boulder beach of probable Pleistocene origin, Iris Point, Orpheus Island.

but some reefs are found near Cape Tribulation (Hopley et al. 2007). Coastal progradation has joined former patch reefs to the mainland at Yule Point and Kurrimine. Fringing reefs are common on the shores of most high islands (>Fig. 21.2.12).

2.4.  Cooktown to Torres Strait Reduced wave energy due to the proximity of the Great Barrier Reef makes this a low-energy coastline characterised by sandy beaches with a narrow zone of low beach

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⊡⊡ Fig. 21.2.11 Hinchinbrook Channel mangroves.

⊡⊡ Fig. 21.2.12 Emmagen Reef, Cape Tribulation. The inner reef was built at a 1-metre higher Holocene sea level.

ridges, about 1 km wide, separated by rocky headlands. The beach ridges are mainly of Holocene age, but fragments of a Pleistocene system may remain separated from the outer series by bare salt pan and saline coastal grasslands. Extensive fringing reefs are associated with all offshore islands, but are not well developed on the mainland.

The most impressive coastal dunes in tropical Australia occur in this region, notably around Newcastle Bay, Orford Bay, Cape Grenville (400 km2), Cape Flattery (700 km2), and Cape Bedford (>Fig. 21.2.13). They consist of elongated parabolic dunes up to 5 km in length and over 100 m in height, many stabilised beneath heath or rainforest, but most of the dune fields formed during glacially lowered

Queensland

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⊡⊡ Fig. 21.2.13 Parabolic dunes near Cape Bedford. (Courtesy Geostudies.)

⊡⊡ Fig. 21.2.14 Cheniers, saline flats, and a mangrove-fringed coast, Princess Charlotte Bay.

sea level episodes from sand sources on the continental shelf (Pye 1982). Intervening lowlands behind coastal embayments show cheniers on mudflats and mangroves fringing creeks and the coastline (>Fig. 21.2.14).

2.5.  The Great Barrier Reef The Great Barrier Reef, the largest coral reef system in the  world, runs for 2,300 km off the northeast coast of Queensland, extending from 24° 07' E northwards to the Gulf of Papua (>Fig. 21.2.15). It is also part of the world’s

largest marine protected area, the Great Barrier Reef Marine Park (339,750 km2), and a World Heritage Area, with a slightly larger area of 347,800 km2. The area of continental shelf within this is 223,997 km2, of which 20,055 km2 or 9% of the shelf area is reef. Including fringing reefs, 2,904 reefs have been recognised. A major review of the evolution of the Great Barrier Reef is contained in Hopley et al. (2007). Fringing reefs occur on the offshore continental islands, but are rare on most of the mainland coast. Many fringing reefs appear to have been substantially in place by the mid-late Holocene, with only minor accretion and/or

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progradation occurring since (Smithers et al. 2006; Hopley et al. 2007). Similar to coral reefs elsewhere, the Great Barrier Reef has been strongly influenced by sea level changes during its evolution (Hopley et al. 2007). During periods of glacial low sea level, the entire continental shelf was exposed, and karst erosion has been responsible for at least some reef features such as “blue holes”, which have been recognised as drowned dolines (>Fig. 21.2.16). Modern reefs came into existence only during the latter part of the post-glacial transgression when the rising

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sea level inundated the older interglacial reefs and allowed recolonisation by corals and other reef organisms. These foundations are generally found at depths between 4 and 12 m below present sea level on the northern and southern Great Barrier Reef, but are deeper, 14–29 m, on the central Great Barrier Reef. Recolonisation took place during the last 10,000 years. Reefs have grown upwards at modal rates of about 7 m in 1,000 years and, because of the differences in depths of the foundations, they have reached a ­post-6,500-year stable sea level at ­different times – about 6,000 years ago from the

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⊡⊡ Fig. 21.2.16 Cockatoo Reef ‘blue hole’ in the Pompey Complex.

shallowest ­foundations, and within the last 2,000 years from deeper platforms.

2.6.  Torres Strait Torres Strait, the shallow 150-km stretch of water between north-eastern Australia and Papua New Guinea, formed a land bridge between the two land masses until as recently as 7,500 years ago, when the final phase of the post-glacial transgression brought the present geography into existence. Located between 9° and 10° S of the equator, the Torres Strait Islands are subjected to a strong seasonality  in climate, with a wet monsoonal period between December and March dominated by light moderate northwesterly winds, contrasting with the remainder of the year, when drier and stronger (25-knot) southeasterlies prevail, though still with possibilities of rainfall. Annual rainfall totals are about 1600 mm. Tropical cyclones at this low latitude are rare, but storms in the Coral Sea can generate large waves that can affect the Torres Strait Islands. A line of large, high islands extends from the Prince of Wales, Horn, and Thursday Island groups near Cape York, through Banks and Mulgrave islands in the central Strait, to Duaun Island near the coast of Papua New Guinea. A further group of high islands is found on the eastern margin of the Strait. They are volcanic in origin, with volcanic activity extending into the Quaternary. The reef

islands are assumed to have formed after the reefs reached sea level, with most beginning to accumulate after about 5,000 years ago, although some, like Warraber, are younger (Woodroffe et al. 2007).

2.7.  The Gulf of Carpentaria The depositional coastline of eastern Gulf of Carpentaria is divided into two parts by the extensive bauxite-rich cliffs of Albatross Bay. From the Jardine River near the tip of Cape York south to Albatross Bay the coastal plain is about 8 km wide, and is characterised by a seaward series of Holocene ridges, a landward degraded pre-Holocene barrier, and an intervening belt of seasonally inundated swampland. North of Vrilya Point a series of beach ridges, derived jointly from the sands of the Jardine River and from the tidal delta at the western exit of Torres Strait, have formed at the coast. Between Vrilya Point and Duifken Point the coastal plain is linear and relatively narrow (2.5–5 km). South of Albatross Bay to the Norman River the coastal plain broadens markedly to a maximum width of nearly 30 km immediately north of the Mitchell River. It is drained by numerous meandering tidal creeks and rivers lined with mangroves. The coastal plain consists largely of low-lying mudflats with beach ridges or occasional low cheniers resting above. The former are more common and usually larger near river mouths, especially to the north,

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where they form a massive cuspate foreland at Cape Keer Weer. Towards the south, stretches of beach ridge coast alternate with chenier plains; this pattern appears to be controlled by the extensive alluvial fans that lie inland of the coastal plain. Narrow beach ridge sequences typically form where the shoreline coincides with the steeper seaward fan foreslopes, whereas cheniers are usually located where broader muddy flats and salt pans have developed in low gradient embayments bounded by adjacent alluvial fans. During the wet season runoff and higher summer tides inundate wide areas of the outer plain. Along the Southern Gulf of Carpentaria (from the Norman River to the Northern Territory border) the coastal plain broadens to form a 40 km-wide chenier plain that is geomorphologically very similar to the chenier plains of the southeastern Gulf of Carpentaria. It consists of discontinuous low, shelly cheniers separated by low-lying mudflats and salt pans, which are inundated by tides and runoff during the wet season. Older Pleistocene ridges lie up to 30 km inland and are sub-parallel to the present coastline. The Wellesley Islands in the southern Gulf consist of laterised Cretaceous and Jurassic sediment, extensively overlain by Pleistocene dune calcarenites rising to 40 m, and modern quartzose beach ridges. Shore platforms are cut into the sediment and fringing reefs surround the islands. The turbid waters of the Gulf of Carpentaria are not known for their coral reefs, but recent research has

discovered up to 80 km2 of patch reefs in water depths of about 28 m. Further multi-beam swath sonar mapping may reveal more extensive reefs in the Gulf.

References Bird ECF, Hopley D (1969) Geomorphological features in a humid tropical sector of the Australian coast. Aust Geogr Stud 7:89–108 Hopley D (1970) The Geomorphology of the Burdekin Delta, North Queensland. Department of Geography, James Cook University, Monograph Series. 1 Hopley D (1971) The origin and significance of North Queensland island spits. Z Geomorphol. 5:371–389 Hopley D, Smithers SG, Parnell KE (2007) The geomorphology of the great barrier reef: development, diversity and change, Vol 532. Cambridge University Press, Cambridge Pringle AW (2000) Evolution of the east Burdekin delta coast, Queensland, Australia: 1980–1995. Z Geomorphol 44:273–304 Pye K (1982) Morphological development of coastal dunes in a humid tropical environment, Cape Bedford and Cape Flattery, North Queensland. Geogr Ann 64A:212–227 Smithers SG, Hopley D, Parnell KE (2006) Fringing and nearshore coral reefs of the great barrier reef: episodic Holocene development and future prospects. J Coastal Res. 22:175–187 Ward WT (2006) Coastal dunes and strandplains in southeast Queensland: sequence and chronology. Aust J Earth Sci. 53:363–373 Woodroffe CD, Samosorn B, Hua Q, Hart DE (2007) Incremental accretion of a sandy reef island over the past 3000 years indicated by component-specific radiocarbon dating. Geophys Res Lett, 34(3) L03602

21.3  Northern Territory    Eric Bird    

1.  Introduction The Northern Territory coastline is about 5,100 km long, and the 148 offshore islands larger than 1 sq. km add a further 2,100 km (Galloway 1985). The eastern part faces the Gulf of Carpentaria, the northern part the Arafura Sea, and the western part Bonaparte Gulf (>Fig. 21.3.1). Bedrock in the eastern part consists of moderately folded, resistant Pre-Cambrian sediment. In the centre, around Darwin, more highly folded ancient sediment,

metamorphic rocks, and granites are found and are overlain by weaker Mesozoic and Cainozoic sandstones and claystones on the peninsulas and islands to the north. In the west, Palaeozoic sandstones and siltstones dip gently westward under patches of Mesozoic sediment. Much of the area has a cover of Cainozoic terrestrial sediment or deep weathering profiles that pass below sea level at many places. The coast is bordered by wide shelves with low relief of subaerial origin thinly mantled by clayey sands and silts inshore. Evidence exists for Pleistocene low sea levels down to −200 m.

⊡⊡ Fig. 21.3.1 Northern Territory: coastal features and mean annual rainfall. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The climate is tropical, with a rainy season from November to April in the north and northwest, where mean annual precipitation is 1,400–1,800 mm. In the southeast and southwest rainfall is less (800–1,100 mm) and the rainy season shorter. Winds in the dry season are southeasterly and generally 10–30 km/h, while in the rainy season weaker westerly and northerly winds predominate, but occasional tropical cyclones bring high winds and intense rains. Mean monthly air temperatures at Darwin range from 25°C in July to 30°C in November: there is a long warm dry winter (April–November) and the prevailing southeast winds are frequently interrupted by a NW sea breeze. Mean spring tide range is 5.5 m at Darwin, about 5.0 m to the southwest, 2.5–3.5 m in Van Diemen Gulf, and 3.5–4.5 m along the north coast. In the Gulf of Carpentaria it is 0.2–1.5 m, but the ranges between highest and lowest tides over a lunar month are 0.5–1.5 m greater. Tidal bores up to several metres high run up the Victoria and Adelaide rivers, while on low coasts storm surges of up to 9 m have been reported. In general wave energy is low, except during tropical cyclones. Currents, partly tidal and partly wind-driven, flow from east to west in the dry season but weaker currents from the west and southwest predominate inshore in the wet season and probably have more influence on sediment transport because rivers are then discharging sediment-laded flood water. Sediment transport in the western Gulf of Carpentaria is from south to north. Water temperatures are around 32°C in the wet season and 27°C in the dry season – high enough for both corals and mangroves (Bardsley et al. 1985). Dune calcarenite, extensive in southern and western Australia, is unusual in the Northern Territory, and its occurrence on Vanderlin Island in the Sir Edward Pellew Group, on Groote Eylandt and on the coast southwest of Cape Arnhem is due to the lack of river sediment and the consequent dominance of biogenic carbonate sediment on the floor of the Gulf of Carpentaria. It was delivered to these coastal regions by southeasterly winds during Interglacial low sea level phases when carbonate sands were exposed on the emerged sea floor, and by wave action during and since the Late Quaternary marine transgression.

2.  The Northern Territory Coastline The Queensland border (along 138°E) crosses the lowlying southern coast of the Gulf of Carpentaria west of Tully Inlet. The coast consists of beaches and beach ridges

backed by low dunes and elongated supratidal flat areas that are submerged as lagoons during and after heavy rains. These are interrupted by the Calvert River and Robinson River deltas. Longshore drifting is westward, and several beaches become longshore spits that deflect tidal creeks in that direction, as at Stockyard Creek and Pelican Spit. The mean spring tide range is about 1.5 m, and at low tide a sandy foreshore is exposed, declining into a shallow nearshore area. There are mangrove fringes to creeks and muddy tidal flats. The coast is backed by a broad plain of Quaternary sediment, but Pre-Cambrian and Cenozoic rocks outcrop in the Sir Edward Pellew Islands and on Maria Island to the northwest (>Fig. 21.3.1). Vanderlin Island, largest of the Sir Edward Pellew Group, has sandy beaches backed by dunes on its southeastern coast. There are fragments of Pleis-tocene dune calcarenite, probably dating from the Last Interglacial, but the parabolic dunes and beach ridges are Holocene. Sandy and muddy areas are exposed at low tide around these islands and on the mainland coast, where McArthur River has built a large delta with branching distributaries in the lee of the Sir Edward Pellew Group. Mangroves are extensive along the tidal channels and along Port McArthur, a narrow strait constricted by delta growth. The pattern of beaches, beach ridges, low dunes and backing supratidal flats continues northwest past the deltas of Rosie Creek and Limmen Bight River, which branches into estuarine distributaries overlooked by Mount Young (65 m) at the northern end of the Yiyintyi Range. In Limmen Bight, the southwestern corner of the Gulf, the meandering Roper River flows through extensive, bare supratidal mudflats seamed by winding, mangrove-fringed tidal creeks. Mangroves fringe river channels, particularly on meander slip-off slopes. Roper River contributes sediment that drifts mainly northward from Port Roper to form sand ridge complexes and mudflats. Despite exposure to waves generated by southeasterly winds, these mudflats support mangroves, although most have been damaged by wave action. They are backed by bare or sparsely vegetated supratidal mudflats which are subject to flood scouring and occasionally interrupted by low clay dunes and windmodified sand ridges (>Fig. 21.3.3). Older and more weathered sandy beach ridges or cheniers and low dune complexes of Last Interglacial age lie farther inland, up to 12 km from the coast. They are similar to features on the east coast of the Gulf of Carpentaria which have been related to sea level a few metres above the present. The east coast of Arnhem Land consists of low, rocky peninsulas and islands, wide mudflats in sheltered bays,

Northern Territory

and extensive sand dunes, behind some of which lie tidal lagoons or lakes. The rocky coasts locally have low cliffs and narrow rock platforms on exposed sites, particularly on soft weathered rock outcrops. There are minor fringing coral reefs, and older emerged coral has been noted within dunes. Mudflats in sheltered gulfs have zoned mangroves up to several kilometres wide, backed by bare supratidal flats that pass inland into freshwater swamps and are drained by straight mangrove-fringed channels. Beach ridges, spits and tombolos, often extensively modified by blowouts, are common on the mainland north to Cape Barrow and in sheltered bays on the islands. Lagoons behind the sand barriers have only modest tidal deltas and narrow mangrove fringes because of the small tidal flux, sandy sediment, and limited riverine input of fresh water and sediment. Sandy beaches fringe the Cape Barrow peninsula. Off Cape Barrow is Bickerton Island, which has sandy  shores and rocky shore platforms on its southern coast. The much larger Groote Eylandt has an embayed eastern and northern coast and a straighter western and southern  coast. Much of the coast has sandy beaches, with beach ridges, cusps, spits that deflect the mouths of creeks, and outlying shoals. Mangroves line sheltered parts of the shore, as in Bartalumba Bay, in the northwest (>Fig.  21.3.2), where they grow behind a fringing coral reef (>Fig.  21.3.3). Parabolic dunes are extensive on the east and south coasts, where they rise to 70 m above sea level. Lake Angurugubia is a brackish lagoon linked to the

⊡⊡ Fig. 21.3.2 Bartalumba Bay, northwest Groote Eylandt.  (Courtesy Geostudies.)

21. 3

sea by a channel where inwashed sand in the dry season is swept out during river floods. Hawksnest Island, northwest of Groote Eylandt, is one of several small islands that have cuspate spits on their lee (northwestern) shores. Blue Mud Bay (>Fig.  21.3.4) is shallow and muddy, bordered by irregular headlands and bays, several of which are fed by small rivers. Wave energy is low because of the shelter of many islands. Intertidal and supratidal mudflats are extensive, with mangroves in the upper intertidal zone initiating swamp encroachment. Numerous parallel creeks, orthogonal to the coastline, indicate relatively rapid and uniform advance of the high tide shoreline. There are few sandy beaches, but some of the headlands have fringing coral reefs. Between Cape Shield and Cape Grey the coast is more exposed, and there are sandy beaches backed by dunes in the bays between rocky headlands. Dunes are spilling northward across the isthmus at Cape Grey. Caledon Bay is a submerged valley system, and to the north the beaches lengthen and are backed by extensive dunes, mainly of parabolic form. The oldest of these may be Pleistocene, but the youngest are modern. Along the coast is a dissected and cliffed Pleistocene dune calcarenite ridge, up to 88 m high, backed by younger parabolic dunes. Much of the calcareous sand was blown in from the emerged sea floor by southeasterly winds during Pleistocene Interglacial phases, but onshore drifting has continued in Holocene times, sand delivered to the beach by wave action being swept inland by southeasterly winds.

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⊡⊡ Fig. 21.3.3 Fringing reef backed by mangroves, northwest Groote Eylandt. (Courtesy Geostudies.)

⊡⊡ Fig. 21.3.4 Blue Mud Bay. (Courtesy Geostudies.)

The dunes are extensive north of Wanyamera Point (>Fig. 21.3.5). Although the climate would sustain luxuriant vegetation there are only scattered patches of dense scrub, and there are numerous blowouts and eroded areas (>Fig. 21.3.6). Grazing by buffaloes has apparently weakened the vegetation cover and accelerated erosion. Port Bradshaw has a narrow entrance and is an intricate mangrove-fringed muddy bay. A long sandy beach

extends northeast to Cape Arnhem, where there are cliffs cut in Pleistocene dune calcarenite (>Fig. 21.3.7). These show pitting and undercutting by solution processes. The northeast coast of Arnhem Land, facing the Arafura Sea, consists of islands, peninsulas, and inlets, the outlines of which are strongly influenced by the geological strike. Rocky promontories and shore platforms alternate with sandy coves and some wider sandy beaches, and there

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⊡⊡ Fig. 21.3.5 Wanyamera Point. (Courtesy Geostudies.)

⊡⊡ Fig. 21.3.6 Dunes south of Binanangui Point. (Courtesy Geostudies.)

are spits that have grown generally eastward, deflecting tidal creeks. Islands and peninsulas have coral reefs on their southeastern shores and small Holocene beach ridges and transverse dunes on the northwest. Sheltered bays are edged by zoned mangroves backed by bare supratidal flats that grade landward into alluvial plains or swamps. Occasional cheniers are found on the flats, which are slightly dissected by straight tidal creeks with broad mangrove fringes. Mangroves also extend far up the rivers and their tributary creeks, and also along form narrow fringes along sheltered rocky coasts. Melville Bay, bordered by the Gove Peninsula, has sandy beaches between rocky promontories. There is a cuspate foreland on the western shore, and a large dunecapped spit has grown southward. A long narrow promontory runs out to Cape Wilberforce, ending in vertical cliffs of stratified and jointed Permo-Carboniferous sandstone over shales, a sequence that dominates much of the coast between here and the Cobourg Peninsula. The sandstones and shales dip gently northwestward, and the outlines of the islands follow this structure. The sequence is repeated in the English Company’s Islands, which have escarpments on their southeastern coasts and embayments and headlands on their northwestern coasts, a pattern clearly seen on Wigram Island. There are structural shore platforms on sandstone bedding planes. Inglis Island, at the SW end of the English Company’s Islands, is also of Permo-Carboniferous sandstone over shale. There are sandy beaches on the north coast of Cape Newbald, and Mallison Island has a 45 m cliff in

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⊡⊡ Fig. 21.3.7 Dune calcarenite, north end of Oyster Bay.  (Courtesy Geostudies.)

­ erm-Carboniferous quartzose sandstone over thinlyP bedded shale, dipping 5°–12° NW. To the south is Arnhem Bay, with mangrove-fringed shores behind intertidal mudflats, interrupted by a few rocky points. On the western coast are cliffs in red Tertiary formations, and there are sandy beaches on Everett Island. Narrow sandstone islands in Ulunourbi Bay converge on Flinders Point. The Flinders Peninsula is also dominated by massive Permo-Carboniferous sandstones, dissected along vertical joints and bedding planes that dip northward at about 5°. The NW coast has numerous small headlands and bays facing Buckingham Bay, a wide bay with a curving southwestern shore which has extensive mangroves and supratidal flats with halophytic vegetation slightly dissected by mangrove-fringed tidal creeks. The curved shore has been shaped by refracted northeasterly waves in a pattern that is  common on sandy beaches and rather unusual on a swampy coast. Inland, there is a transition to freshwater swamps that show evidence of former tidal channels. The southeastern coast of Napier Peninsula consists of rugged hummocky cliffs cut in horizontal and gently undulating white and grey sandstones, with many rocky lobate headlands between sandy coves. The Wessel Islands are another chain of long narrow islands running out NE into the Arafura Sea, on PermoCarboniferous sandstones dipping NW at up to 20°. Their SE coasts are steep escarpments, their NW coasts embayed with fringing reefs, particularly around promontories. There are abrupt cliffs on Point Napier, at the NE end of

the Napier Peninsula. Strong tidal currents sweep through Stretton Strait northeast of Elcho Island. This island has long sandy beaches and cliffy headlands with fringing coral reefs on its northern coast (as at Naningbura Point), exposed to northwesterly wave action from the Arafura Sea, and swamp-fringed Cadell Strait to the south, passing SW into a mangrove-fringed tidal channel behind Howard Island. Banyan Island, on the southern coast of Castlereagh Bay, also has cliffs and sandy beaches on its northern side, Yabooma Island has a cuspate foreland at its eastern end, and Milingimbi is largely fringed by mangroves and intertidal mudflats, in front of mangrove-fringed islands and creeks. The coast runs out to a sandstone promontory at Cape Stewart. In the more open situation of Boucaut Bay, sandy beach ridges and low parallel dunes are found behind the shore, especially in the eastern part, exposed to stronger northwesterly winds and waves. The sand ridges are recurved and splayed on either side of the Blyth River estuary. Mangroves are confined to narrow tidal flats in their lee, and supratidal flats have occasional cheniers. A spit has grown eastward, deflecting Anamayirra Creek, and to the west the coast becomes more sheltered and swampy. Farther west, at the mangrove-fringed mouth of Liverpool River there are sandy coves between rocky points where ironstone and laterite cap pink and red flatlying sandstone at Maningrida. A series of sandstone headlands and islands, with swampy bays at river mouths, extends northwest to Junction Bay, which has similarities to Boucaut Bay, and

Northern Territory

on west of Cuthbert Point. The Goulburn Islands to the west have cliffs cut in soft weathered rocks, beach ridges, and low transverse dunes on the exposed northwestern shores and spits and small parabolic dunes on the southeastern shores, related to dry season southeasterly winds. Beaches fringe much of the coast that runs northwest to Cape Cockburn. On the northern coast of Cobourg Peninsula and Croker Island, there are elongated bays formed from wide, partially drowned valleys, with intertidal flats that bear extensive mangroves (Wells 1982). Bays exposed to the northwest have extensive sand accumulations including spits, transverse dunes, and beach ridges. Smaller islands on platform reefs have prominent ridges of coral sand. Fringing reefs up to 3 km wide run round headlands and are backed by mangroves or sand accumulations. Cape Don is attached to the mainland by a mangrove isthmus. Dundas Strait, the northern entrance to Van Diemen Gulf, is broad and up to 140 m deep, to the east of Cape Don, but passes southward into splayed ridges and channels having a local relief up to 30 m: they are tidally shaped features and may include some drowned reefs. The low southern coast of Cobourg Peninsula, facing the eastern part of Van Diemen Gulf, is exposed to winds and waves from the southeast. Steep cliffs in soft weathered rock are retreating rapidly but are soon degraded where the growth of beach ridges at their feet has checked erosion. Patchy Holocene parabolic dunes extend 1–2 km inland, and dunes are also climbing degraded cliffs.

⊡⊡ Fig. 21.3.8 Field Island, the Cunningham Channel and Midnight Point, southeastern Van Diemen Gulf. (Courtesy Geostudies.)

21. 3

The eastern and southern coasts of Cobourg Peninsula are muddy, with extensive mangroves that have been damaged by storms unless protected by fringing coral reefs. In the southeast the mangroves are protected by beaches and sand ridges but are narrow and backed by supratidal flats with occasional cheniers and sedge-covered alluvial plains (>Fig.  21.3.8). Isobaths suggest drowned barriers exist offshore. On Field Island mangroves lie in front of beaches, a common situation from this point westward along the southern coast of Van Diemen Gulf. The West Alligator River flows out through a channel bordered by mangroves, which extend back along creeks that drain supratidal swamps (>Fig. 21.3.9). The northern coasts of Melville Island and Bathurst Island are exposed to northwesterly waves across the Timor Sea, and have cliffed headlands between long gently curving sandy beaches backed by dunes. Brenton Bay, Lethbridge Bay, Snake Bay and Shark Bay are long estuarine inlets that formed by partial submergence of river valleys and have been maintained by strong tidal currents, the tide range here being about 6 m. Currents are very strong through Apsley Strait, the narrow channel that separates the two islands. On the northern and northwestern coasts spits have formed, partly blocking the mouths of bays and inlets. A major intertidal sand shoal extends out northwest from Cape Van Diemen, shaped by the interaction of strong tidal currents with Timor Sea waves. The southern coast of Melville Island, facing Clarence Strait, is less exposed, and has forested slopes descending

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Northern Territory

⊡⊡ Fig. 21.3.9 The mouth of the West Alligator River. (Courtesy Geostudies.)

to high tide level, with coral reefs fringing Cape Gambier. The intertidal zone is broad, partly sandy and partly muddy, and there are elongated shoals and reefs between east-west channels shaped by the strong tidal currents in Clarence Strait. The southern coast of Bathurst Island facing Beagle Gulf has an eastern part which is protected by spits, sand ridges, and fringing reefs, behind which lie mangroves, bare supratidal flats, cheniers, and degraded, scalloped cliffs. To the west it swings WSW and is exposed to strong southeasterly winds and waves. Cliffs up to 20 m high have been cut in weathered rock, but they are now vegetated, and fronted by a wide beach, probably as the result of Holocene emergence. Active parabolic dunes occupy low-lying sectors. The smooth southern coast of Van Diemen Gulf has an almost continuous fringe of mangroves up to 500 m wide, backed by bare supratidal flats that grade inland into dark grey clay plains that are sedge-covered or bare of vegetation and cracking in the dry season. Cheniers, strips of sand generally about 100 m wide, deposited on the clay plains, are marked out by lines of Pandanus trees (>Fig. 21.3.10). Behind Finke Bay there are multiple cheniers (>Fig.  21.3.11) which converge westward into the cuspate promontory of Point Stuart, diverging again further west. Point Stuart has a shelly beach (>Fig. 21.3.12). Clarke et al. (1979) have shown that the cheniers at Point Stuart are of Holocene age, and that relative sea level has not been more than 1 m above its present position during the last 6,000–7,000 years. At Sampan Creek a chenier has been washed inland, leaving an intertidal wave-cut slope on swamp clay to seaward.

The cheniers extend up to 18 km inland, and consist of quartzose sand with some ironstone nodules and much shell (but no coral) sand: the shells are mud-dwelling species that have drifted in from the broad mudflats that extend below low tide level in Van Diemen Gulf. They were probably storm-built, perhaps emplaced during occasional tropical cyclones. They are intersected by straight or slightly meandering mangrove-fringed tidal creeks that indicate rapid and uniform progradation. There has also been vertical accretion, the fringes of the older (inner) cheniers being overlapped by wedges of blue clay. The hinterland is an extensive confluent river floodplain, across which the South Alligator River, the Mary River and the Adelaide River flow to funnel-shaped tidal estuaries (Woodroffe et al. 1986, 1993). Narrow fringes of mangrove extend alongside the main rivers up to 80 km inland, nearly to the tidal limit. The nearshore waters are shallow and muddy. The western entrance to Van Diemen Gulf is largely blocked by reefs, notably the oval platforms of the Vernon Islands, which are up to 15 km long and support extensive mangroves. Partially drowned laterite surfaces may also be present. Fringing reefs around headlands, Greenhill Island, and other islands provide protected sites for mangroves that grow up to 15 m high. Lee Point, a rocky promontory with a boulder beach and ferruginous beach rock with coral fragments, is where the coast swings south towards Darwin. Casuarina Beach, wide at low tide, has swash bars and rippled swales in fine, firm sand strewn with crab pellets at low tide. It is backed by low foredunes held by Spinifex, breached by the outlet

Northern Territory

21. 3

⊡⊡ Fig. 21.3.10 Pandanus trees on sandy chenier, coastal plains south of Van Diemen Gulf. (Courtesy Geostudies.)

⊡⊡ Fig. 21.3.11 Multiple cheniers, Finke Bay, Van Diemen Gulf. (Courtesy Geostudies.)

of a tidal creek that has been deflected northward. To the south Dripstone has low cliffs with caves, clefts and ledges cut out along joints and bedding-planes in the almost horizontally stratified Lower Cretaceous limestones and sandstones of the Mullaman Group, capped by clays and ironstone gravels (>Fig. 21.3.13). Small headlands correspond with minor anticlines. A mangrovefringed creek opens to the southern end of the beach, where cliffs with structural ledges run out at Nightcliff.

Erosion has exposed a fault line scarp that runs across the shore, and the rock outcrops are pitted and honeycombed by weathering. Barbed wire and other metallic waste from Second World War defences has become coagulated into a rusting ironstone. South of Nightcliff mangroves fringe a muddy bay at Bagot, and sandstone cliffs resume on East Point. On the southern side the cliff descends to a sloping abrasion ramp, formed where waves have moved coarse sand and gravel

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Northern Territory

⊡⊡ Fig. 21.3.12 Shelly beach ridges, Point Stuart, Van Diemen Gulf. (Courtesy Geostudies.)

⊡⊡ Fig. 21.3.13 Cliffs and cave in Cretaceous sandstone at Dripstone, Darwin, the sandstone overlain by Tertiary clay and ironstone g ­ ravel. (Courtesy Geostudies.)

to and fro on the upper shore. In Fannie Bay, to the south, the sandy beach includes coralline gravel, and erosion has exposed layers of beach rock, sloping gently seaward and disintegrating (>Fig. 21.3.14). Beach rock also occupies joints and fissures in the Cretaceous sandstones (>Fig. 21.3.15). The horizontal Cretaceous sandstone and limestones form shore ledges and are seen to rest upon the Pre-Cambrian basement, exposed at Bullocky Point in the southern cliffs. Darwin is built on a promontory bordered

by cliffs and bluffs with bouldery shores and segments of cemented beach conglomerate. Port Darwin has a branching configuration, with five marine inlets that formed as the result of partial marine submergence of former river valleys, and are now lined with extensive mangroves. The promontories between the valleys include low cliffs and shore platforms cut in weathered rocks, beaches with beach rock, beach ridges and small transverse dunes formed by summer northwesterlies.

Northern Territory

21. 3

⊡⊡ Fig. 21.3.14 Disintegrating beach rock on the shore at Fannie Bay, Darwin. (Courtesy Geostudies.)

⊡⊡ Fig. 21.3.15 Beach rock (calcareous cemented sandstone) in joint fissures, Fannie Bay, Darwin. (Courtesy Geostudies.)

West of Darwin there are long marine and estuarine inlets between narrow peninsulas that end in sandstone headlands such as Charles Point on Cox Peninsula. There are sandy beaches and segments of fringing coral reef. There are spits on Masson Point, Rankin Point and Point Celyon. Thrings Channel leading into Bynoe Harbour is geomorphologically similar to the Darwin district and Port Patterson is another long branching mangrovefringed macrotidal inlet. On its western side coral reefs

link a chain of islands north to Quail Island. To the south multiple sandy beach ridges extend to the mouth of Finnis River, beyond which is a mangrove-fringed swamp. Fog Bay is more exposed, facing west, and has long sandy beaches shaped by waves from the Timor Sea. On the cliffs at Point Blaze there is an emerged beach of coralline sand and gravel 3 m above high tide level. South from Point Blaze low cliffs and shore platforms on weathered rock outcrops alternate with broad bay beaches having beach ridges

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Northern Territory

at their more exposed northern ends, mangroves behind sand spits and barrier island in the centre, and mangroves at their more sheltered southern ends. Behind the bays lie extensive swampy supratidal flats, which may have been intertidal prior to Holocene emergence. The swam­ ­py  supratidal flats become bare of vegetation as rainfall decreases southward. In Anson Bay the Daly River has built up extensive alluvial clay plains, mostly above high tide, but with narrow mangrove fringes at the coast and low cheniers on the western side of the estuary. Isobaths indicate a shallow submarine delta. Daly River sediment are carried northward and contribute sand and mud to the Peron Islands, the northern one being high and ridged, the southern low and mangrove-fringed. Cliff Head to the south is cut in Tertiary sediment, and on Red Cliff these overlie PermoCarboniferous sandstone, which emerges here and dominates Cape Ford and cliffs to the south. South of rocky Cape Ford more sandy beaches alternate with cliffed headlands and shore platforms with some fringing coral reefs. Cape Scott is a tombolo, an island attached to the mainland by a sandy isthmus with numerous beach ridges, and to the south there are multiple beach ridges on either side of a tidal inlet with a mangrove swamp. There are cliffs and structural shore platforms cut in red, white and yellow stratified Permo-Carboniferous sandstone with a capping of Tertiary sand south to Cape Dombey. Moyle River drains a broad swampy plain to flow into Hyland Bay, and White Cliff Point is the beginning of a line of cliffs which run out to Tree Point, a long rocky peninsula with cliffs cut in Permo-Carboniferous rocks bordering Port Keats, another mangrove-fringed drowned valley system. A spit that has grown northeastward from Munda Beach has almost blocked the tidal channel south of Dorcherty Island, which ends in cliffs of Permo-Carboniferous sandstone. Munda Beach is backed by multiple beach and dune ridges, and cliffs and bluffs alternate with sandy beaches along the coast to Pearce Point, another cliffed headland of Permo-Carboniferous sandstone, which marks the beginning of the broad macrotidal Joseph Bonaparte Gulf. Within Joseph Bonaparte Gulf coastal waters are usually clouded with fine suspended sediment from the rivers, and there are many rippled shoals, mostly NNW-SSE, awash at high tide. The main channels near the coast are shallow and partially blocked by rapidly changing muddy

shoals and islands readily colonized by mangroves. The coast consists of mangrove-fringed swamps and supratidal mudflats. Fitzmaurice River cuts through sandstone ridges and widens to Keyling Inlet, while Victoria River has a funnel-shaped estuary opening to Queens Channel and Keep River flows out through a similar widening bay beside Turtle Point. Each of these rivers has a macrotidal estuary, and delivers a large seasonal sediment supply. Extensive intertidal and supratidal flats between the main channels are threaded by tidal creeks. Mangroves fringe the rivers and creeks, but the supratidal flats are vast bare surfaces occasionally submerged by the highest tides or river floods. The tidal creeks and mangroves may be extending headward into these bare flats. On the open coast mangroves have suffered storm damage, even though the shallow water attenuates the waves. The Victoria estuary penetrates 100 km inland, cutting through steep ranges of quartzose sandstone with escarpments facing southwest, and upstream are the broad Whirlwind Plains, where muddy sediment is derived from shales. At Whirlpool Reach the ebb tide attains 15 km/h, and during spring tides has a tidal bore runs through Gunns Reach and far upstream. The river channel is bordered by a flood plain that is occasionally inundated by exceptionally high tides and in places supports stunted mangroves. West of the Keep River estuary the wide coastal plain continues across the border (along 129°E) into Western Australia.

References Bardsley KN, Davie JDS, Woodroffe CD (1985) Coasts and tidal wetlands of the Australian monsoon region. Australian National University, Northern Australian Research Unit, Mangrove Monograph 1 Clarke ME, Wasson RJ, Williams MAJ (1979) Point Stuart chenier and Holocene sea levels in northern Australia. Search 10:90–92 Galloway RW (1985) Northern Territory. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold. Stroudsburg, ­Pennsylvania, pp 949–956 Wells AG (1982) Mangrove vegetation of northern Australia. In: Clough BF (ed) Mangrove ecosystems in Australia. Australian National University Press, Canberra, pp 57–78 Woodroffe CD, Chappell JMA, Thom BG, Wallensky E (1986) Geomorphological dynamics and evolution of the South Alligator Tidal River and Plains, Northern Territory. Australian National University, Northern Australian Research Unit, Monograph 2 Woodroffe CD, Mulrennan ME (1993) Geomorphology of the Lower Mary River Plains, Northern territory. Northern Australian Research Unit, Darwin

21.4  Western Australia    Ian Eliot . Eric Bird       

1.  Introduction Western Australia covers an area of 2.5 million sq. km, extending from latitude 13° 30' S to 35° 08' S and from longitude 113° 09' E to 129° E. The 12,500 km coastline comprises one-third of the Australian total, and includes a wide range of coastal environments. Western Australia has entirely different climates in its northern and southern parts, while in the central region there is a gradual change from the tropical climate of the north to the Mediterranean climate of the south. Weather systems are largely controlled by movement of an anticyclonic wind belt, which brings easterly winds to the northern half of the state and westerlies to the south. Prevailing wind patterns in the north are disrupted by late summer tropical cyclones, while in the south mid-latitude depressions occur during winter (Gentilli 1971). Landforms on the coast reflect the dominant wind, wave, current, and tidal processes and their interaction with the varied basement geology. Western Australia is largely made up of Pre-Cambrian igneous, metamorphic, and sedimentary rock. The remainder of the state consists of younger sedimentary rocks, mainly sandstone and limestone, that flank the older rocks (Geological Survey of Western Australia 1975). In places, sedimentary rocks extend offshore as a wide continental shelf. Coastal landforms of the Kimberley and South Coast are strongly controlled by the geological structure of resistant basement rocks. Elsewhere, structural control is weaker, and wind, wave, and tidal processes determine gross coastal configuration. Coastal processes vary with the prevailing and dominant weather conditions, coastal aspect, and changes in the tidal regime. A southwesterly swell generated in the Southern Ocean affects most of the coast. As a result, the south coast is subject to a moderate-to-heavy southwesterly swell; the west coast to a moderate southwesterly swell; and the north coast to a low, refracted westerly swell (Radok 1976). Waves generated by tropical cyclones, inter-anticyclonic fronts and depressions, mid-latitude ­depressions, and sea breezes are superimposed on the prevailing swell, so that the wave regime varies markedly

for different parts of the coast. In general, littoral currents transport sediment from the southwest corner of the state north towards Carnarvon and east towards Esperance. On the macrotidal northwest coast sediment transport is dominated by tidal currents, with wave action playing a subordinate role, but the rest of the coastline of Western Australia is characterised by wavedominated zeta-form embayments of various sizes. North of Cape Leeuwin these face either north or northeast, whereas east of Cape Leeuwin they generally face southeast. Eight coastal regions are delineated on the basis of geological structure, dominant processes, and landforms (>Fig. 21.4.1). These are described in turn from north to south.

2.  The Kimberley Coast The Kimberley coast is here defined as extending from the Northern Territory border west to King Sound and Cape Leveque. It fringes a region largely composed of resistant Pre-Cambrian (Proterozoic) rocks (about 1,800 million years old) that include crystalline granite, metamorphic schist and quartzite, and volcanic lavas. The topography is strongly influenced by strong southeast-northwest and secondary northeast-southwest trending joints and fractures. Coastal processes are characterised by a prevailing low westerly swell, southeast winds, and very large tide ranges. Semidiurnal spring tides generally exceed 7 m, and attain a maximum of 10.5 m in Collier Bay. The intertidal zone is correspondingly wide, and very strong currents are generated as the tides rise and fall. The climate is tropical, with a summer wet season (December–March) and a dry winter when southeast trade winds are dominant. Tropical cyclones, which generally move in a southwest direction, parallel to the coast, periodically generate strong onshore winds and high wave conditions, notably in the wet summer season. Coastal configuration is determined by the structural grain of the hinterland. The west coast of Joseph Bonaparte Gulf, between the Northern Territory border and Cape Londonderry, is linear and parallel to the major

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Western Australia

⊡⊡ Fig. 21.4.1 Coastal regions of Western Australia. (Courtesy Geostudies.)

southeast-northwest joint pattern, except where the wide Cambridge Gulf (a structural depression) opens at the mouth of the Ord River. This funnel-shaped estuary is bordered by extensive supratidal saline flats, with alluvial plains at the mouth of the Ord River, constituting a large delta. Long linear shoals (Medusa Banks and King Shoals) diverge between deeper channels offshore. At Cape Londonderry the coast swings westward. Cape Londonderry and Cape Talbot are fringed by broad

coral reefs. The coastline is irregular, with many rias. Rivers debouch into broad gulfs and long, narrow inlets, bordered by cliffs up to 75 m high in Pre-Cambrian rocks (Kimberley Group). Sandstones are extensive, and Cape Bougainville is cut in bauxitic laterite. There has been marine submergence of deeply-incised bays, bordered by high peninsulas with precipitous rocky headlands, and by numerous islands (Bonaparte Archipelago) and reefs. The influence of southeast-northwest fractures is

Western Australia

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⊡⊡ Fig. 21.4.2 Tidal waterfall, Kimberley coast. (Courtesy John Marquet.)

seen in the long straight valley of Prince Regent River, opening into St. George Basin. Within the framework of steep, rocky coasts there are shoals and depositional features. Sediment distribution in the region is dominated by onshore-offshore tidal currents and periodic storm wave activity, resulting in the formation of discontinuous tidal flats, cheniers and deltas, and mangrove swamps (Thom et al. 1975). Mangroves occur on sheltered shores, as at the mouth of Roe River in Prince Frederick Harbour and at the head of Doubtful Bay. The large tides generate tidal bores in narrow gulfs, and racing upstream in the rivers. South of Walcott Inlet the Pre-Cambrian rocks of the King Leopold Range have swirling outcrops that run out on an intricate peninsula of rias and cliffed ridges, fringed by the rocky islands of the Buccaneer Archipelago. There are local coastal outcrops of granite and dolerite. In this region the large rising and falling tides are locally confined between high rocky islands to form tidal waterfalls (>Fig. 21.4.2). These have sometimes been called “horizontal falls”, but this term is erroneous. South of the Buccaneer Archipelago is King Sound, a major drowned valley system at the mouth of Fitzroy River. The mean spring tide range is more than 8.0 m, increasing to 8.9 m at Derby, at the southern end, where maximum tides of up to 12 m have been recorded. The Sound is up to 60 km wide at high tide, with elongated peninsulas in the south and numerous reef-fringed islands of the Buccaneer

Archipelago in the north. Extensive saline flats, submerged only at high spring tides, storm surges or wet season flooding, border much of eastern and southern coast of the Sound, with a serrated inner coastline related to protruding scrubby dune ridges (red pindan sand) and muddy inlets along intervening swales. The pindan ridges are part of the Great Sandy Desert, and are thought to have formed during an arid phase about 20,000 years ago, in the late Pleistocene, when the sea level was lower, King Sound was a broad valley, and the coastline of Dampier Peninsula lay far to seaward of its present position. On the west coast of King Sound there are several broad, branched embayments of the kind that are termed sebkhas on the coasts of the Red Sea and the Persian Gulf. They are often partly enclosed by sandy spits and barriers, and penetrated by the sea to an extent that varies with tides and weather. Above the normal level of marine submergence there are salt-encrusted areas, invaded by the sea only during the higher tides and storm surges. They branch into inter-dune swales and show parallel zones of different evaporite deposits (chlorides, sulphates, and ­carbonates in a landward sequence), formed by precipitation from seawater and from fresh water springs around their margins. Chattur Bay on the west coast of King Sound is a good example (>Fig. 21.4.3). The west coast of King Sound has cliffy sectors, cut in horizontally bedded Cretaceous sandstones and shales. These are prominent at Cape Leveque (>Fig. 21.4.4),

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⊡⊡ Fig. 21.4.3 Chattur Bay, a sebkha on the west coast of King Sound. (Courtesy Geostudies.)

⊡⊡ Fig. 21.4.4 Cliff cut in ferruginous sandstone, Cape Leveque. (Courtesy Geostudies.)

and extend down the west coast of Dampier Peninsula to  Broome and beyond. Cliffed promontories separate sandy beaches, backed by dunes. Offshore strong tidal  currents sweep between the many rocky islands and reefs of the Buccaneer Archipelago as the tides ebb and flow.

3.  The Canning Coast The Canning Basin extends inland from the coast between Cape Leveque and Port Hedland. It includes the Dampier Peninsula, the Broome Peninsula, Roebuck Bay, and the De Grey delta.

Western Australia

Cape Leveque is the northernmost point of the Dampier Peninsula, a triangular area of semi-arid pindan scrub on red sandy and silty soils (an aeolian brickearth, also termed pindan). The surface is almost featureless, except for shallow valleys that open westward to tidal estuaries in Pender Bay and Beagle Bay. These estuaries are fringed by mangroves, backed by saline flats that are submerged only at the highest tides, and an inner shoreline with fresh water seepages. A segmented coastal ridge of Pleistocene dune calcarenite forms a series of headlands. Further south is the long sandy beach, backed by dunes, that becomes Cable Beach at Broome. At low tide the sandy shore is wide and gently sloping, with faint parallel bars, swale lagoons, and seepage outflows. Locally, there are blowouts in the dunes, exposing ridges and hummocks of dune calcarenite. At the southern end of the beach a backshore ridge of cobbles has been derived from the rocky outcrops and boulders on the shore curving out to Gantheaume Point. The cliffs on Gantheaume Point show structural ledges on horizontal red and grey Cretaceous Broome Sandstone strata (>Fig. 21.4.5). Dinosaur footprints can be seen at low tide. Pindan-capped sandstone cliffs extend southwest to Entrance Point, where the Port of Broome jetty runs out. The mean spring tide range at Broome is 8.2 m. Roebuck Bay is about 25 km wide, facing west. Man­ groves fringe its inner shores (>Fig. 21.4.6). It has a gently sloping sea floor of sand and mud, and at low spring tides the shore is up to 2 km wide, increasing to 8 km in the southern part of the bay. Cretaceous sandstones outcrop ⊡⊡ Fig. 21.4.5 Gantheaume Point. (Courtesy Geostudies.)

21. 4

in low cliffs and rocky foreshores, with stacks and natural arches (>Fig.  21.4.7), and pass laterally into cliffs of red pindan (>Fig. 21.4.8). From Bush Point the coast runs southwest, and successive sand spits that have grown to the northeast are interspersed with muddy tidal flats and mangroves bordering Yardoogarra Creek. Cape Villaret is the first of a series of low, cliffy promontories that separate sandy bays. The wide beaches are often fronted by orthogonal sand bars extending seaward and partly exposed at low tide. The sequence of promontories and beaches comes to an end at Cape Missiessy, which marks the start of the Eighty Mile Beach. This long sandy beach, backed by pale dunes, curves southwest, then west, and is backed by salt flats and marshes, followed by an irregular inner shoreline, the limit reached by the highest tides and summer flooding. The beach sand is partly quartzose and partly calcareous. It is backed by several parallel dune ridges that narrow to the southwest and overlie abandoned Pleistocene tidal flats. Inland from Mandora is a re-entrant in the pindan dunes where a chain of salt lakes marks a former estuary. A residual salt creek is fringed by an unusual inland occurrence of mangroves (Avicennia marina), described by Beard (1967) and thought to be relict from a higher sea level. To the southwest the Eighty Mile Beach merges with the De Grey River delta, a major lobate structure extending about 20 km beyond the general trend of the coast. This is the only instance in Western Australia where fluvial deposition occurs on such a scale on the open coast.

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⊡⊡ Fig. 21.4.6 Mangrove fringe at Broome, on the shores of Roebuck Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 21.4.7 Natural arch in Cretaceous sandstones on the shores of Roebuck Bay. (Courtesy Geostudies.)

4.  The Pilbara Coast The Pilbara coast extends from Port Hedland to Cape Preston. It borders the Pilbara Block and the ­Hammersley

Basin, composed of Archean and Proterozoic rocks, including granite, metasediment, and volcanic rocks. Coastal processes are similar to those for the Kimberley region. This is the impact zone for the majority of tropical

Western Australia

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⊡⊡ Fig. 21.4.8 Cliffs cut in pindan (brickearth) on the shores of Roebuck Bay. (Courtesy Geostudies.)

cyclones crossing the Western Australian coast. Strong storm surges, high wave activity, and exceptional heavy rainfall in the inland catchments that feed the De Grey, Fortescue, Ashburton, and Gascoyne rivers are associated with these events. Otherwise, the coast is affected by westterly sea breezes and a weak, refracted northwest swell. Basement rocks outcrop as rocky headlands and islands of the Dampier Archipelago. Broad tidal flats, cheniers, and storm ridges form the coast within the embayments. The eastern embayment is characterised by an exceptionally broad band of tidal flats and a chain of elongate islands composed of sandy calcarenite. The islands are remnants of a large zeta-form Pleistocene embayment. Holocene tidal flats overlie extensive relict Pleistocene tidal formations landward of the islands.

5.  The Carnarvon Coast The Carnarvon coast extends from Cape Preston to Kalbarri, and borders the Carnarvon Basin. Along the coast, sediment of the Carnarvon Basin are horizontally bedded in the north, folded into large north-northeast trending anticlines in the west, and uplifted on the southern margin where they abut the Pre-Cambrian Northampton Block. In conformity with the change of aspect at North West Cape, the dominant coastal processes vary southward as the open coast becomes subject

to the prevailing south-southwesterly Southern Ocean swells and a strong southerly wind wave regime. East of the cape the sheltered coast is subject to a refracted northwest swell and weak offshore winds. The spring tide range diminishes from 5 m in the north to about 2 m at Carnarvon. The eastern shore of Exmouth Gulf has a broad expanse of tidal flats and fringing mangroves. South of the cape the coastline consists of high cliffs cut into folded Tertiary rocks. A Pleistocene coral reef occurs at the base of the cliffs. Offshore, a modern fringing barrier reef, Ningaloo Reef, extends over 300 km along the coast, from False Island Point to Pont Quobba. The Gascoyne River forms a lobate delta that lies immediately north of an abandoned Pleistocene delta. Pleistocene limestones were formed at the extremity of a northerly sediment transport system that operated along the west coast of Western Australia during the Pleistocene. Spectacular cliffs, ranging up to 100 m in height, characterise the exposed coastline. Shark Bay is a major region of carbonate sedimentation and for the modern occurrence of stromatolites in the hypersaline waters of Hamlin Pool (Logan et al. 1974).

6.  The West Coast The West coast, from Kalbarri to Cape Naturaliste, borders the Perth Basin, in which Phanerozoic sediment

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provide a basement to Quaternary coastal sand dune sequences that include lithified Pleistocene calcarenities and unconsolidated Holocene calcareous sands (Semeniuk and Meagher 1981). Onshore and offshore ridges of limestone run subparallel to the present coast, outcropping as dunes, chains of reefs, and offshore islands. Behind gaps in these reefs are Holocene sediment in complex beach ridge sequences forming cuspate forelands and tombolos, as in Jurien Bay. Pleistocene and Holocene dune sequences incorporate estuaries and extensive, elongate en echelon coastal lagoons and lakes. South to Fremantle low cliffed bluffs with extensive rock platforms and reefs alternate with sandy bays and beaches. On more exposed parts of the coast, sporadic outcrops of beach rock occur (Russell and McIntire 1965). Holocene sands frequently form large parabolic dune sequences and transgressive sand sheets overlying similar massive Pleistocene landforms. South of Fremantle limestone outcrops occasionally interrupt long sandy beaches in zeta-form embayments. The dominant headland is Cape Naturaliste, which shelters Geographe Bay. On a smaller scale, a columnar basalt headland at Bunbury and a limestone headland at Mandurah control smaller, shallow embayments. The sandy beaches alternate seasonally between a low-energy, summer condition with cusps to a higher-energy, rhythmic morphology with rip currents in winter. Other features of the coast include the coral reefs and islands of the Abrolhos, a major Holocene sediment

accumulation at Rockingham, and suites of rock platforms on Point Peron and Rottnest Island (Fairbridge 1950). These, and the extensive dune sequences of the Swan Coastal Plain, have formed during changes in Quaternary sea level.

7.  The Leeuwin Coast Near Bunbury there is an isolated stand of mangroves (Avicennia marina) (>Fig. 21.4.9) on the shores of Leschenault Inlet (33° 20' S, 115° 38' E) (Semeniuk and Withers 2000). The nearest mangroves are about 600 km to the north in Houtman Abrolhos and about 1,800 km to the east on the coast of the Eyre Peninsula in South Australia. On the Cape Naturaliste peninsula Archaean granite and gneiss are overlain by Pleistocene dune calcarenite and calcareous dune sand. The west-facing coast south from Cape Naturaliste to Cape Leeuwin is generally steeply sloping, with some cliffed sectors and rocky shores on granite-gneiss where the Archaean basement outcrops (>Fig. 21.4.10), and on calcrete where Pleistocene dune calcarenite extends down to or below sea level. There are sandy surf beaches on which the dominant longshore drifting is northward. The Leeuwin Naturaliste National Park is of varying width, extending 100 km southward along the coastal range, which declines eastward into a gently undulating plain extending from Geographe Bay in

⊡⊡ Fig. 21.4.9 The southernmost mangroves in Western Australia at Mangrove Park, Leschenault Inlet, Bunbury. (Courtesy Geostudies.)

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⊡⊡ Fig. 21.4.10 Domed outcrops of granite gneiss near Cape Leeuwin. (Courtesy Geostudies.)

the north to Flinders Bay at Augusta in the south. Margaret River has cut a valley running southwest through the coast range, and the river mouth is encumbered by sand shoals. Cape Leeuwin (34° 22' S, 115° 08' E) is a low granitegneiss promontory capped by thin patches of dune calcarenite. This is the most southwesterly point on the Australian mainland, where the Indian Ocean meets the Southern Ocean.

8.  The South Coast East of Cape Leeuwin an irregular granite-gneiss coastline runs to Flinders Bay, where the low, rocky coast turns northward to a sandy cuspate foreland with low grassy dunes beside the entrance to Hardy Inlet, a large estuarine lagoon at the mouth of Blackwood River. It shows only minor fluctuations in level (tidal range up to 0.5 m), except when the sea is raised or lowered by meteorological effects. On the eastern side a beach-fringed coastline curves smoothly for 35 km round to Black Point, a cliffed promontory of hexagonal basalt. A series of headlands, sandy bays, and river mouths extends to the western side of the bold D’Entrecasteaux promontory. Point D’Entrecasteaux rises 122 m above sea level. The bordering cliffs are of dune calcarenite capped by grey calcrete, with a steep slope in orange and yellow dune sand and fallen calcrete blocks. The outlying Flat (or Quangering) Island consists of a granitic reef with a low

grassy dune, and Sandy Island to the east has grassy dunes on a similar hard reef. Windy Harbour is a sandy bay to the east of Point D’Entrecasteaux, with a beach interrupted by many outcrops of hard metamorphic rock. A large, sandy cusp on the shore has formed as the result of accretion in the wave convergence zone behind Rat Rock. The coastal dune fringe is backed landward by a broad swampy corridor that runs east-southeast to Brok, a brackish lagoon about 30 km long and up to 4 km wide, linked to the sea by a narrow channel through the coastal dune fringe. Mandalay Beach is a high, wide sandy beach backed by a sector of calcarenite cliff that includes dark brown humate (>Fig. 21.4.11), a sandstone bound by organic matter, possibly formed in a lake or swamp in the accumulating Pleistocene dunes. Walpole stands on the northern shores of Walpole Inlet, a lagoon less than a metre deep, which receives Walpole River. It is linked by a strait southward through a hilly ridge to estuarine Nornalup Inlet, which is up to 5 m deep. Successive promontories of Archaean rock separate bays with curved, sandy beaches and river mouths that are often shoaly. An embayed cliffy coastline runs eastward past Rame Head, which has a vertical cliff cut out along a joint plane in dark dolerite. At Point Irwin the coast turns northward and declines to Peaceful Bay, a cove with several rocky outcrops. Madfish Bay, lies in the lee of a low, elongated ­granite-gneiss island, Madfish Island. This island is capped by tors and boulders, with small dunes of white

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⊡⊡ Fig. 21.4.11 Dark humate outcrop in the dune calcarenite cliff near the western end of Mandalay Beach. (Courtesy Geostudies.)

sand bearing stunted scrub. The intervening shallow strait is constricted by a cuspate sand spit that extends out from the mainland shore, shaped by converging waves in the lee of the island. To the east is Waterfall Bay, a sandy cove backed by a low humate cliff, over which a stream falls 1.5 m. The humate was formed in the quartzose dunes by subsoil precipitation of iron oxides and organic matter leached from the overlying sand by podzolic weathering. Wilson Inlet leads into a large coastal lagoon south of Denmark. The entrance channel is constricted on the eastern side by a large dune-capped sand spit descending to a sandy hook, which impacts currents against the dune calcarenite cliffs on the western shore. A granite-gneiss coast emerges near Sharp Point. Torndirrup National Park is a heathy granitic landscape, with tors and some quartzose sand dunes over dune calcarenites. Locally, the calcarenite mantle has been stripped off to expose domed granitic surfaces which have been dissected along vertical joints to form several chasms (geos), one of which, the Natural Bridge, is spanned by a slab of granite (>Fig. 21.4.12), while another nearby is The Gap, a chasm 24 m deep. Along this coast the Pre-Cambrian (Archaean) rock outcrops show evidence of multiple intrusions: banded gneiss with parallel wavy lines of light and dark minerals formed 1,350 million years ago, intruded by coarsegrained grey or pink granite with xenoliths (inclusions) of gneiss and black dolerite seams, formed about 1,180

million years ago. These rock types are matched on the Antarctic coast near Windmill Island, indicating the separa­­tion  of Australia from Antarctica in Cretaceous times; the south coast of Western Australia is still moving northward. The granite-gneiss coast runs out in Flinders Peninsula to Bald Head. On the northern side is King George Sound, a large bay with Michaelmas Island and Breaksea Island, high islands that shelter it from the east. Beaches and low granitic outcrops run behind Frenchman Bay on its southern coast, and the port of Albany on its northern side. The normal tidal range here is up to 0.6 m, and tides are largely diurnal. As on other parts of the south coast of Western Australia, fluctuations of sea level resulting from changes in barometric pressure or alternations of onshore and offshore winds can raise or lower the water level ­independently of, and on a larger scale than, normal tidal oscillations. Albany stands on the southern slopes of a granitegneiss ridge (Mount Melville, Mount Clarence, and Mount Adelaide) that runs east to King Point on the coast of King George Sound. To the east of the Emu Point Strait are rocky granite-gneiss shores. The coast east from Beaufort Inlet is backed by high dunes, largely covered in scrub and woodland, but with many blowouts. These are backed by a wetland corridor. On the east coast of the Point Henry Peninsula, Short Beach is backed by low dunes. The backshore is not eroded, perhaps because a chain of nearshore parallel dark reefs

Western Australia

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⊡⊡ Fig. 21.4.12 Natural Bridge on the coast south of Albany. People standing on the bridge indicate the scale. (Courtesy Geostudies.)

⊡⊡ Fig. 21.4.13 Waves breaking over an intertidal calcarenite reef west of Esperance. (Courtesy Geostudies.)

protect it. There are only gentle waves on this eastern shore, compared with the much larger swell and surf on coasts with a southwest aspect. The shores are rocky rather than cliffed. The Bremer Bay breakwater is a large structure of schist blocks quarried nearby. The adjacent rocky shore is of sandstone and infolded schist. The beach curves round to Jones Cove, a promontory, beyond which is Peppermint Beach on the Hood Point Peninsula. The white sand hereabouts is calcareous, not quartzose.

To the north is the Fitzgerald River National Park, with a generally steep coast north of Point Ann on the west coast of Doubtful Island Bay, and a sandy beach curves past the outlet from Culham Inlet. This had been generally sealed by a dune-capped barrier until 1993, when an outlet reopened during a flood. From the mouth of Culham Inlet curving beaches and minor points extend to West Bay, and on to Flathead Point, a cuspate spit in the lee of nearshore reefs. Seaweed and seagrass are often heaped on the beach at Hopetoun, where

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a breakwater running out from a small promontory shelters the harbour from easterly waves. To the east of Hopetoun a scrubby dune bluff and backshore dune terrace are fronted by a white sandy beach. A cuspate point runs out to a small group of brown schist rocks. At Two Mile Beach a steep dune bluff descends to a dune calcarenite cliff behind a narrow, sandy beach with a planed-off calcarenite ledge above high tide level, then a blue lagoon and waves breaking over an intertidal calcarenite reef. Between Hopetoun and Esperance the coast consists of sandy bays with some calcarenite, separated by small rocky headlands of Proterozoic migmatite. As elsewhere on the West Australian coast, the dunes and dune calcarenite formerly extended further seaward, and have been cut back during Holocene times to leave nearshore segments of intertidal and subtidal calcarenite reef (>Fig. 21.4.13). Stokes Inlet at the mouth of Young River is an estuarine lagoon bordered by outcrops of schist, sectors of shelly sand, and swamp scrub. To the east is rocky Shoal Cape, then a series of small asymmetrical bays with sandy beaches shaped by ocean swell refracted round intervening headlands. Towards Esperance, Eleven Mile Beach is backed by a slope with low wind-pruned heath, and on Ten Mile Beach intertidal calcarenite segments en­­ close  a lagoon, except for a gap at Ten Mile Lagoon (>Figs. 21.4.14 and >21.4.15). The town of Esperance is similar to Portland in >Victoria, especially the orientation of the coast and the harbour breakwater at Dempster Head. North of the

­ reakwater beaches have been depleted by erosion, which, b as at Portland, may be due to modification of incident wave patterns by the breakwater. Esperance Bay is a sandy beach backed by high scrubby dunes, then a lagoon bordered by salt marsh with dead paper barks and rushes, and sectors of salted scrub in a former wetland. Cape Le Grand National Park occupies a broad promontory on Pre-Cambrian granite and gneiss. At Thistle Cove steep smooth coastal slopes of pink granite are incised by clefts cut along joints. To the east is Rossiter Bay, with a long, curving sandy beach that is often heaped with dry seagrass (Zostera). Behind Rossiter Bay is a series of headlands and bays, bordered by the many islands of the Archipelago of the Recherche, which extends eastward to Cape Arid on dune calcarenite and Cape Pasley, where Tertiary limestone caps Proterozoic metamorphic rocks. The coast then swings northeastward past Point Malcolm, and becomes more arid.

9.  The Eucla Coast The Eucla coast extends from Cape Arid to the South Australian border, fringing the Eucla Basin, which is occupied by Miocene limestone that forms a flat plateau 50–60 m above sea level. This is the Nullarbor Plain, a remarkable plateau, essentially the depositional surface of the Miocene Nullarbor Limestone, a sedimentary plain uplifted and tilted gently southward and eastward, preserved without any river valleys, although there are ⊡⊡ Fig. 21.4.14 An emerged calcarenite bench on Blue Haven Beach, Esperance, showing disintegration along the undercut seaward margin. (Courtesy Geostudies.)

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⊡⊡ Fig. 21.4.15 An air view of Esperance, showing the harbour breakwater, the Pink Lake, and Lake Warden. (Courtesy Geostudies.)

underground caves (Jennings 1963). The long, almost straight cliffed coastline along its southern margin is probably not faulted: it is the outcome of marine erosion of a coast of uniform height and uniform rock resistance. Beyond Point Malcolm the Wylie Scarp forms the southern margin of the Nullarbor plateau (Lowry 1971). The scarp is fronted by a low-lying coastal plain, with a Holocene barrier system in the vicinity of Israelite Bay, enclosing lagoons that have become salt flats, such as Lake Daningdella, which has evaporite deposits, including halite and gypsum. Mean annual rainfall in the coastal fringe is about 225 mm, but in this dry and windy environment evaporation is strong. The spectacular vertical Baxter Cliffs occur where the Wylie Scarp comes to the coast between Point Culver and Twilight Cove. The cliffs then pass inland as a scarp (a relict cliff line, formed when sea level was higher) bordering the Hampton Tableland (Nullarbor Plain). This Hampton Range scarp is fronted by the Roe Plains, an emerged lowland up to 40 km wide, between Eyre and Eucla. The scarp has an upper cliff of Nullarbor Limestone above a sloping apron on the less resistant Wilson Bluff Limestone. To the east this sloping apron becomes basally cliffed. The basal cliff then in-creases in height, so that the middle slope fades out and the cliff disappears beneath the Merdayerrah Sandpatch. When the cliff reappears the sloping apron revives, until it passes beneath the Eucla dunes at the Head of the Bight. The arid Roe Plains carry a Pleistocene dune ­calcarenite sequence, overlain by Holocene dunes along the ­seaward

margin and extending below present sea level. Wind, wave, and tide regimes in this region are similar to those of the South Coast. Winds are dominantly from southern quadrants, and strong southwesterly swell is refracted around Cape Arid. Wave action weakens as the continental shelf widens east of Cape Pasley, but storm waves occasionally attack the coast. Towards Eucla an extensive mobile sand sheet with active barchan dunes covers much of the plain, and the Delisser Sandhills occupy the coastal fringe. The scarp comes to the coast at Wilson Bluff, near the South Australian border (129° E), as cliffs in white chalky Eocene limestone, the Wilson Bluff Limestone, overlain by the Nullarbor Limestone.

References Beard JS (1967) An inland occurrence of mangrove. West Australian Nat 10:112–115 Fairbridge RW (1950) The geology and geomorphology of Point Peron, Western Australia. J R Soc West Aust 34(3):35–72 Gentilli J (ed) (1971) Climates of Australia and New Zealand. Elsevier, Amsterdam, the Netherlands Geological Survey of Western Australia (1975) The geology of Western Australia. Memoir 2 Jennings JN (1963) Some geomorphological problems of the Nullarbor Plain. Trans R Soc S Aust 87:41–62 Logan BW, Read JF, Hagan GM, Hoffman P, Brown RG, Woods PJ, Gebelein CD (1974) Evolution and diagenesis of quaternary carbonate sequences, Shark Bay, Western Australia. American Association of Petroleum Geologists Memoir 22

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Lowry DC (1971) Explanatory notes on the Eucla-Noonaera geological sheet Western Australia. Australian Bureau of Mineral Resources, Canberra Radok R (1976) Australian coasts. Rigby, Perth Russell RJ, McIntire WG (1965) Southern hemisphere beach rock, Geogr Rev 55:17–45 Semeniuk V, Meagher T (1981) Calcrete in quaternary coastal dunes in southwestern Australia: a capillary-rise phenomenon associated with plants. J Sediment Petrol 51:47–67

Semeniuk V, Withers P (2000) The Leschenault inlet estuary. J RSoc West Aust 83 Thom BG, Wright LD, Coleman JM (1975) Mangrove ecology and deltaic-estuarine geomorphology: Cambridge Gulf-Ord River, Western Australia. J Ecol 63:203–232

21.5  South Australia

C. R. Twidale · J. A. Bourne

1.  Introduction Although some 3,700 km long, the South Australian coast is in broad view simple in plan, with only two major indentations and one large island interrupting and separating the NW–SE curve of the western coast and the offset NW–SE trending southeastern sectors (> Fig. 21.5.1). The coast was explored in late 1801 and early 1802 by

Matthew Flinders, working eastwards, and by Nicolas Baudin travelling to the west. Despite their two countries being at war at the time, the two explorers met and amicably exchanged information in the well-named Encounter Bay. Many of the present names of coastal features, such as Spilsby Island, Boston Bay, and Cape Catastrophe on the one hand, and Murat Bay, Cape Carnot, and Cape du Couedic on the other, are due to these early explorers.

⊡⊡ Fig. 21.5.1 Location map and tectonic regions.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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2.  Tectonic Elements Elements of four tectonic terranes a terrain is a crustal blocks or fragment that preserves a distinctive geologic history, different from the surrounding areas and usually bounded by faults are exposed on the South Australian coast (> Fig. 21.5.1). From west to east, they are the Eucla Basin, the Gawler Craton including the Stuart Shelf; the Adelaide Geosyncline, an orogen of Early Palae­ozoic age with which the Torrens Graben is associated; and in the southeast, the Padthaway Horst, on to which lap the Late Cainozoic strata of the Gambier Basin (the western part of the more extensive Otway Basin). This last region has also witnessed Late Quaternary hotspot volcanism. The NW–SE trends of the west coast of Eyre Peninsula and of the Upper and Lower South East districts parallel lineaments, in these instances, complex fault zones, as do the east coast of Eyre Peninsula, the west coast of Yorke Peninsula, and the northern margin of Kangaroo Island.

3.  The Eucla Coast The Eucla coast, the seaward margin of the Eucla Basin, lies between the Western Australian border and the Head of Bight and consists mainly of 50–60 m high cliffs which rise sheer from the Southern Ocean (> Fig. 21.5.2). They are

developed in white Nullarbor Limestone of Miocene age, overlying beds of the Eocene yellow-brown Wilson Bluff Limestone. Known as the Bunda Cliffs, there are locally developed narrow shore platforms. Blocky talus accumulations mark zones of recent cliff recession. Near Eucla, at the western end of the Bunda Cliffs, a platform cut in ­calcarenite and standing some 5–6 m above sea level fronts the cliffs and extends westwards into the Roe Plain (> Fig. 21.5.3), where the backing scarp is erosional, and not tectonic as has been surmised. The Wilson Bluff Limestone that is exposed in this locality includes lenses and pods of flint which were gathered for tool-making by the indigenous people. Climbing and cliff-top dunes occur at several sites (Jennings 1967). The Bunda Cliffs offer no direct evidence of sea level change, but the horizontal zonation of the deeper caves below the Nullarbor Plain suggests they are graded to a sea level 40–50 m higher than present.

4.  The ‘West Coast’ From the Head of Bight to the southern extremity of Eyre Peninsula (> Fig. 21.5.1), the coastline is developed on plutonic, metamorphic, and sedimentary rocks of the Gawler Craton which are patchily exposed beneath a cover of Quaternary dune sediment and most com­mon­­ly dune ­calcarenite. In this region, known in the State of South

⊡⊡ Fig. 21.5.2 Bunda Cliffs, with brown Miocene overlying white Eocene limestone. Note the extreme flatness of the Nullarbor etch plain. (Courtesy Coast Protection Board, South Australia.)

South Australia

21.5

⊡⊡ Fig. 21.5.3 Lower platform in calcarenite standing 70–80 m below the Nullarbor Plain and some 8–10 m above sea level. The unconformity between the Late Pleistocene calcarenite and the Tertiary limestones is exposed in the cliff in the foreground. Note climbing dunes in distance. (Courtesy Coast Protection Board, South Australia.)

Australia as the ‘West Coast’, the ­calcarenite is of MiddleLate Pleistocene age, with dates of 630,000–180,000 years recorded from thermoluminescence deter­minations.

5.  Old Foredunes Calcarenite dunes extend 90–100 km inland from the present coast on the essentially stable western Eyre Peninsula and are also preserved on such islands as the Investigator Group, where old foredunes link the North, Middle, and South sections of Pearson Island. Flinders Island is capped by calcarenite resting on plutonic rocks, of which granite gneiss is exposed on the coast, though the presence of a kimberlite pipe is suspected and the Island is an active diamond prospect. On Emu and several other small offshore islands, however, only calcarenite is exposed. As can be seen in the cliff sections, the sequence comprises stacks of calcarenite dunes with palaeosol horizons, commonly calcrete. Unconformities, and particularly, the calcrete horizons within the dune sequence, give rise to prominent ledges (> Fig.  21.5.4). The calcrete displays a nodular habit (‘golf-ball’ structure) in places and solution pipes (orgues géologiques) preserved by secondary calcification elsewhere. The calcarenite is commonly 150 m or so thick and on Thistle Island sections more than 200 m thick are exposed. The dunes developed at a time of lower sea level, for the calcarenite, extends below the present low tide level at

several sites (−19 m at Elliston), but the unconformity between the dune rock and the underlying granitic materials and sediment is irregular. Where it is exposed, a regolith (ferruginised on granitic bedrock) is preserved. The dune topography consists of rolling hills devoid of surface streams, and on the coast, this rolling character finds expression in the varied heights of the cliffs. Typically, the land slopes gently inland from the cliff line. Numerous dolines and short caves are developed and aligned hilltop dolines have been attributed to underprinting. The coastal dunes have isolated former coastal embayments in lagoons such as lakes Hamilton and Newland. On the other hand, Baird, Coffin Bay and its extensions in Port Douglas, Mt Dutton and Kellidie bays, Streaky, Venus, and Waterloo bays all maintain narrow openings to the sea. The field of coastal dunes also blocked the Eocene and Pliocene Narlaby drainage system which previously emerged to the sea at what is now referred to as Smoky Bay, where the former river course is marked by a channel. Waterloo Bay is almost circular in plan and of particular interest, for it has been identified as a possible meteorite impact crater. No evidence of such an origin has been found, however, and the almost circular bay may be either due to the coastal barrier being breached and wave refraction, then producing the present regular outline, or it may be an especially large doline. The former is more likely, but in both instances, development could have been influenced by fractures underprinted from the underlying granite basement.

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⊡⊡ Fig. 21.5.4 Cliffs in old calcarenite dunes, west coast of Eyre Peninsula. (Courtesy Coast Protection Board, South Australia.)

6.  Bedrock Forms The coastline consists of alternations of rocky headlands fronted by stacks and shore platforms, and beaches backed by modern dune fields. In detail, the character of the headlands varies according to structure for though all are shaped in dune limestone, the morphology of the associated platforms varies according to the level of the unconformity with respect to sea level, and thus, to the nature of the rock exposed. Smooth shore platforms have been shaped in calcarenite by marine weathering and wave erosion. The platforms are narrow on promontories due to deep quarrying by waves, but they attain considerable width in bays and on sheltered coasts. In some instances, their presence is more frequent, and their extent greater, which may be obvious at first sight, for some are covered by beach sand for most or all of the year. Remnants of ­platforms standing just above the usual high tide level and higher platforms of limited extent are present at many sites. Wave ramps also are in evidence. Flats occur with­­in  the spray zone and are attributed to water-level weathering. These platforms developed in Middle-Late Pleistocene calcarenite are related to the present stand of the sea which, give or take one metre, attained its present level 6000–7000 years ago. On the west coast of Eyre Peninsula platforms extend up to 300 m and commonly 200–250 m from the base of low limestone cliffs, suggesting an average rate of cliff recession and concurrent platform extension of 3–4.5 cm/year. Sea caves (and blowholes), arches

and stacks, as well as solution cups and flutings, are characteristic of this calcarenite coast. On the granitic islands of the Investigator archipelago sheet structure is prominent, but elsewhere the morphology of exposures of Precambrian rocks varies partly with fracture density but also according to the elevation of the unconformity. At Point Labatt and Smooth Pool, for example, the unconformity between calcarenite and granite stands within the present tidal zone and the granite occurs in smooth platforms. At Point Drummond, on the other hand, the unconformity stands almost 20 m above sea level and a ferruginised sub-calcarenite regolith is well exposed. At Point Sinclair, the calcarenite rests on a partly exhumed bornhardt which rises some 15 m above sea level. Alveolar weathering is developed at several sites, though less commonly and well on granitic rocks than on basic intrusions.

7.  The Talia Coast On a short sector of coast near Talia, a gently tilted Precambrian sandstone and grit is exposed. Here the platform is irregular and consists of a miniature cuesta landscape. Prominent fractures disposed at an obtuse angle to the coast have been weathered and eroded to form geos, and in places have concentrated dissolution sufficiently to influence karst development in the overlying calcarenite: hence, The Tub, a doline some 20 m diameter and 10 m deep, bottoming on the sandstone. Sandstone exposed in

South Australia

cliffs below the calcarenite display alveolar weathering,  and an assemblage of rock doughnuts and fonts (> Fig. 21.5.5) has been attributed to contrasts in rates of weathering on exposed rock surfaces and surfaces covered by shell-grit beach rock which holds moisture (Twidale and Campbell 1998). It has been suggested that such weathering may have a wider significance. In particular, such sub-beach weathering may have contributed to the formation of shore platforms (Twidale et al. 2005).

7.1.  Beach Etching Many platforms on the South Australian coast, whether shaped in calcarenite, dolerite, gneiss, granite, or quartzite, are covered by sand and are exposed in whole or in part only after storm waves have combed down the beach. Many carry patches of sand or gravel (shingle). The unconsolidated sediment retains water so that even at low tide and in drought, weathering continues at the base of the beach. Some platforms like that developed at Point Drummond (Molina Ballesteros et al. 1995) are classical etch forms, for remnants of the regolith are preserved in adjacent promontories and cliffs and the block- and boulder-strewn platforms are exposed weathering fronts. Beaches are ephemeral, but the retained water has the same function as shallow groundwaters, and many platforms, such as those occasionally and partly exposed at Greenly Beach and at Back Beach, west coast of Eyre Peninsula, that

⊡⊡ Fig. 21.5.5 Fonts in cross-bedded sandstone, Talia, with calcarenite cliffs in the background.

21.5

revealed to seaward of the beach at Point Riley, Yorke Peninsula, at Hallett Cove on the Gulf coast south of Adelaide, and at Beachport in the South East district, all appear partly to owe their planation to beach etching.

8.  Depositional Forms Some beaches occur in bayheads. Others on low coasts are several kilometres long and are backed by active dunefields. The modern dunes are predominantly white, but behind the northern Talia beaches they are a dun color, hence, Mt. Camel, the highest point in this complex. The dunes display well-developed mobile crests which ‘smoke’ in high winds. The beaches are predominantly calcareous and consist of the fragmented shells of molluscs and other marine organisms. The only exceptions occur where granite is exposed just offshore, as at Tyringa Beach, but even here there is enough shelly material in the beach to be winnowed and accumulated by the wind in foredunes behind the beach. At many sites the beach sand is reinforced by beds of the sea grass (Posidonia spp.) washed up during storms and giving rise in places to low (ca. 1 m) cliffs. In addition, storm waves erode cliffs in the beach sand while at some sites, e.g., at Edward Bay, cusps are frequently formed as to be semi-permanent features. Spits and bars with mangroves (Avicennia marina) occur in sheltered bays and inlets, as for instance at Thevenard and the mouth of Streaky Bay.

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South Australia

On Pearson Island terrace deposits, the outer edge of which stands some 6 m above present sea level, attest a higher stand of the sea, as does the shingle beach preserved below the later Pleistocene calcarenite and exposed in a cliff west of the quartzitic hill known as The Frenchman, on southern Eyre Peninsula. This deposit, like those noted elsewhere (q.v.), may offer opportunities for dating using luminescence techniques.

9.  The Gulf Coasts Spencer Gulf and Gulf St. Vincent occupy depressed fault blocks overridden by the Gawler Craton to the west and the Adelaide Geosyncline to the east (> Fig. 21.5.1). Landforms similar to those described from the west coast of Eyre ⊡⊡ Fig. 21.5.6 Gneissic granitic rocks in serrated platform exposed beneath beach sand at Point Riley, just north of Wallaroo, northwestern Yorke Peninsula. (Courtesy Coast Protection Board, South Australia.)

Peninsula are developed but are more subdued, save at the toe of Yorke Peninsula where there are high precipitous calcarenite cliffs. As on the Eyre Peninsula coasts, Precambrian granites and gneisses are exposed in places beneath the calcarenite. Here, bouldery platforms and quite extensive pool weathering platforms are developed, as are rock platforms exposed from beneath beach deposits (> Fig. 21.5.6).

10.  Old Foredunes The age of the calcarenite, as determined by magneto­ stratigraphy, is similar to Eyre Peninsula occurrences, while a thermoluminescence date of 125,000 years was obtained for the base of a dune exposed in coastal cliffs on northwestern Yorke Peninsula (D. Beng, pers. comm. June 2000). Mostly, however, cliffs shaped in weakly lithified Cainozoic sediment are lower. Erratics are exposed in coastal sections where Permian glacigene strata are eroded as at Port Vincent. Shore platforms are fewer, though those in exposed igneous and metamorphic rocks are irregular and block-strewn.

11.  Constructional Forms Both older and younger coastal dunes occur along these sheltered coasts, but they are lower and less extensive than on the west coast of Eyre Peninsula. Pleistocene desert dunes preserved by calcrete cores extend below sea level both on northwestern Yorke Peninsula and at Lucky Bay, northeast of Franklin Harbour, on the Gulf coast of Eyre Peninsula. Constructional forms such as spits, bars and cuspate forelands, many with mangroves, are common. For example, Franklin Harbour is an embayment with mangrove-clad shores partially enclosed by a spit which has extended northeastwards. Cuspate forelands such as Sultana Point and Black Point are common also on the east coast of Yorke Peninsula, and Weeroona ‘Island’, near Port Pirie, is connected to the mainland by a tombolo. The distributary channels of the Broughton River have formed a delta covered by mangroves.

12.  Higher Sea Level Stands A shingle beach is preserved beneath the dune calcarenite at Hillock Point near the toe of Yorke Peninsula and shingle beaches stand some 3 m above mean sea level on both western and eastern shores of upper Spencer Gulf (Hails et al. 1984). A shingle ridge 6000–4000 years old has been

South Australia

identified at several sites on the eastern shore of northern Spencer Gulf (Burne 1982).

13.  Coasts of the Adelaide Geosyncline The coasts of Kangaroo Island and Fleurieu Peninsula (> Fig.  21.5.1) can be considered as one because both regions are part of the Adelaide Geosyncline, a Palaeozoic fold belt subject to recurrent and continuing earth movements. Tectonics and structure find expression both in gross and in detail. Backstairs Passage, which separates Kangaroo Island from the mainland, may be due to the exploitation of a tensional zone in the country rock, at various times, by rivers, glaciers, and the sea.

14.  Kangaroo Island On Kangaroo Island spectacular cliffs, serrated platforms and sea arches are eroded in the folded Cambrian strata. Permian glaciated pavements are exposed near Pen­neshaw and elsewhere. At Cape Cassini salt weathering has produced notable alveolar weathering in flat-lying sandstones. To the east, the Bay of Shoals is well-named, and Ante­ chamber Bay is backed by a long sandy beach as well as high tide or storm level platforms eroded in calcarenite. Granitic exposures give rise to cliffed headlands, many of which ­display blocks in situ and corestone boulders, as at Cape

⊡⊡ Fig. 21.5.7 Boulder beach on Windmill Bay, eastern Kangaroo Island. Storm beach boulders, only occasionally washed by waves, distinguished by reddish algal growth. Cape Willoughby lighthouse on skyline.

21.5

Willoughby. Nearby beaches consist of boulders which are released corestones (> Fig. 21.5.7), many of which are fitted. Granitic headlands are present on the south coast, most notably at Remarkable Rocks, or Kirkpatrick Point. The feature comprises a domical hill surmounted by angular blocks with well-developed tafoni and flutings. Adja­cent platforms are typically rough and strewn with blocks and boulders (some with flared flanks, others with Kluft­karren). Other­ wise, the south coast of the Island is dominated by sandy beaches and backing dunes, behind which are impounded shallow lagoons, such as Lake Ada and Murray Lagoon. Shells from near the base of the calcarenite cliffs at Point Reynolds, on the south coast, have been dated by the Strontium isotope method which gave an age of 520,000 years, comparable to the older western Eyre Peninsula sequence. Apart from those at the western end of the Island, the calcarenite cliffs are lower than in western parts of the State. On Kangaroo Island the dunes have an added and greater influence on the configuration of the coast, for the narrow neck of land north of Pennington Bay and including Mt Thisby, standing 101 m above sea level, consists of coastal dunes which link what had previously been two islands (> Fig. 21.5.1). On the other hand, during periods of lower Pleistocene sea levels large areas of the present continental shelves were exposed. The Gulfs and Investigator Strait became riverine plains, and the then River Murray extended some 60 km to the SSW before tumbling over the edge of the continental slope where its turbid waters eroded ­submarine

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South Australia

canyons some 1.5 km deep (Hill et al. 2005). It also incised up to 60 m into a granitic landscape which provided the sandy load that facilitated the formation of the turbidity currents responsible for the canyons. High stands of the sea are evidenced both on Kangaroo Island and Fleurieu Peninsula by abandoned stacks, as at Windmill Bay, and by pre-calcarenite dune beaches, like that exposed in Admirals Arch at Cape du Couedic (which is comparable with those mentioned from The Frenchman and Hillock Point). Widespread evidence of stands of the sea some 8 m and 40–44 m above present sea level takes the form of raised beaches and perched coastal river terraces.

15.  Fleurieu Peninsula On the opposite side of Backstairs Passage Fleurieu Peninsula is in most respects similar to the Palaeozoic bedrock coasts of Kangaroo Island. Where sedimentary strata are exposed, the headland and bayhead beach morphology reflects contrasts in rock type and disposition of strata. Encounter Bay is a literal as well as a littoral inselberg landscape with granite islands and promontories, one with a remnant of a glaciated pavement, rising above a plain in metamorphic rocks now partly inundated by the  sea. At the seaward margin of the Willunga Basin south of Adelaide, Early Tertiary beds give rise to smooth shore platforms (> Fig.  21.5.8) backed by steep ledged cliffs: the  steeply-dipping Proterozoic argillites exposed at  Hallett  Cove near Adelaide have been eroded into a

serrated platform backed by structurally–controlled cliffs (> Fig. 21.5.9), and like many other platforms is protected by a rampart on its seaward edge. On the crest of the backing cliff is a well-known Late Palaeozoic glaciated pavement and glacial erratics, some of granite others of the Neoproterozoic Sturtian tillite in glacigene deposits nearby. Like that at Edward Bay, the adjacent beach frequently displays cusps (> Fig. 21.5.10). On the east coast of Gulf St. Vincent, beach drift is from south to north. This is responsible for the northward extension of the Le Fevre Peninsula across the diverted mouth of the Port River. Most of the spit or beach ridge was constructed by waves and wind between 7,500 and 5,000 years ago (Bowman and Harvey 1986). Presumably this reflected the ready availability of sand at the then recently established sea level. It is for this reason also that the maintenance of Adelaide’s metropolitan beaches involves the transportation of drifted sediment from Semaphore in the north back to Marino and Brighton in the south; to date, by trucking the sand, but piping a sand slurry is planned. North of Le Fevre Peninsula the shore consists of mangroves, mud and sand flats. Gulf St Vincent is sheltered and rates of cliff erosion and platform extension of about 3 cm/year are indicated in argillitic outcrops. ­

16.  The South East District To the east of the gulfs, in the South East district, tectonism has disturbed the basement rocks of Palaeozoic ⊡⊡ Fig. 21.5.8 Cliffs and platform in flat-lying Late Eocene Blanche Point marl south of Adelaide.

South Australia

21.5

⊡⊡ Fig. 21.5.9 Serrated platform in steeply dipping Proterozoic siltstones backed by cliffs which are largely exposed bedding planes, Hallett Cove, south of Adelaide. (Courtesy G.W.C. Lyne.)

⊡⊡ Fig. 21.5.10 Beach cusps at Hallett Cove.

age and with them the coastal foredunes developed during  sea level changes dating back almost 800,000 years (Cook et al. 1977). The southeastern area has been uplifted,  while the northwestern, around the Murray Mouth, has subsided (> Fig. 21.5.11). The result is a complex series of subparallel calcarenite dunes, some of which have been active more than once, but which broadly speaking decrease in age from the inland, around

Naracoorte, to the present coast (Sprigg 1952). Remnants of the pre-dune landscape are exposed in various small bouldery granitic hills. The Upper South East district is occupied by a single active coastal beach-dune barrier known as the Youn­ghus­ band Peninsula, behind which is impounded The Coorong (> Fig. 21.5.11), an elongate salt-water lagoon well-known for  its modern dolomite formation. Where the Murray

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South Australia

⊡⊡ Fig. 21.5.11 The Murray Mouth, Younghusband Peninsula, with active foredunes and The Coorong seen from the northwest. (Courtesy Resource Information Group, DENR.)

approach­es the sea, it flows into lakes Albert and Alexan­ drina and the foredune barrier is breached by the Murray Mouth. Southeast of Kingston, where granite boulders are exposed on the beach, the Lower South East district coast consists of calcarenite cliffs of moderate height fronted by shore platforms and with arches and stacks also developed (> Fig. 21.5.12). Lagoons and lakes such as Bonney,

George, and Eliza are impounded behind barrier beaches and associated foredunes. The beaches of the South East ­district display Posidonia-reinforced cliffs, sand cliffs, and cusps. Other depositional forms include tombolos and cuspate forelands. Pollution has killed seagrass allowing mobilisation of sand by waves and resulting in the translocation of about 100,000 m3 of sand per annum off Brighton (a coastal

South Australia

21.5

⊡⊡ Fig. 21.5.12 Low calcarenite cliffs with shore platforms, stack, and arch at Robe. (Courtesy Coast Protection Board, South Australia.)

suburb of Adelaide) alone and the lowering of the seafloor by a metre or more. The varied morphology of the South Australian coast is due to the interplay of structure and marine process­ ­es,  active at various and changing levels, through time. The impacts of Quaternary events are particularly prominent, for the phases of dune-building on west- or southwest-facing coasts have left an enduring imprint. Marine processes, both erosional and depositional, continue to shape the coast at surprisingly rapid rates. Notwithstanding major advances over the past several decades, however, many aspects of the coast remain enigmatic. It continues to pose enough problems and to offer possibilities sufficiently enticing to retain its fascination for the geologists, geomorphologists, and environmentalists alike.

References Bowman G, Harvey N (1986) Geomorphic evolution of a Holocene beach-ridge complex, Le Fevre Peninsula, South Australia. J Coastal Res 2:35–362

Burne RV (1982) Relative fall of Holocene sea level and coastal progradation, northeastern Spencer Gulf, South Australia. BMR J Aust Geol Geophys 7:35–45 Cook PJ, Colwell JB, Firman JB, Lindsay JM, Schwebel DA, Von Der Borch CC (1977) The late Cainozoic sequence of southeast South Australia and Pleistocene sea-level changes. BMR J Aust Geol Geophys 2:81–88 Hails JR, Belperio AP, Gostin VA (1984) Quaternary sea levels, northern Spencer Gulf. Australia. Mar Geol 61:373–389 Hill PJ, De Deckker P, Exon NF (2005) Geomorphology and evolution of the gigantic Murray canyons on the Australian southern margin. Aust J Earth Sci 52:117–136 Jennings JN (1967) Cliff-top dunes. Aust Geogrl Stud 5:40–49 Molina Ballesteros E, Campbell EM, Bourne JA, Twidale CR (1995) Character and interpretation of the regolith exposed at Point Drummond, west coast of Eyre Peninsula, South Australia. Trans R Soc South Aust 119:83–88 Sprigg RC (1952) The geology of the South-East province, South Australia, with special reference to quaternary coastline migrations and modern beach developments. Geol Surv South Aust Bull 29 Twidale CR, Bourne JA, Vidal Romani JR (2005) Beach etching and shore platforms. Geomorphology 67:47–61 Twidale CR, Campbell EM (1998) Development of a basin, doughnut and font assemblage on a sandstone coast, western Eyre Peninsula, South Australia. J Coastal Res 14:1385–1394

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21.6 Victoria Introduction

Eric Bird

1.  Introduction

2.  The Coastline

This account of the coastline of the State of Victoria is presented in the following nine chapters:

The coastal outlines of Victoria have been much influenced by these tectonic movements. The uplifted areas of the Otway Ranges, the Mornington Peninsula and the South Gippsland Highlands each have cliffed coastal margins, and the sea has occupied the subsided basins, Port Phillip and Westernport Bays and Corner Inlet. The Gippsland Lakes lie within a former marine embayment in the structural trough that extends eastward from the Latrobe Valley and out under Bass Strait. West of the Otway Ranges, the Tertiary rocks dip beneath Newer Volcanics, which form cliffs and rocky shores on the Portland Peninsula; eastward they decline into the Port Phillip basin. In the far east of the state the coastal cliffs are cut across the folded and faulted rocks of the Eastern Highlands of Australia, which here show north-south trend. Wilsons Promontory stands out to the south as a Devonian granitic intrusion which has been stripped of its sedimentary roof, then deeply dissected to form an embayed mountainous peninsula south of Corner Inlet. To the west, Cape Liptrap is a cliffed promontory cut into an uplifted ridge of Lower Palaeozoic formations, which include Cambrian rocks - the oldest seen on the Victorian coast. The coastal geology is shown in (>Fig. 21.6.1). The outlines of the uplifted and subsided coastal sectors have been modified during Pleistocene and Holocene times by the extensive deposition of sandy beaches, dunes and barriers (i.e. beaches and dunes in front of lagoons and swamps) along the coast, and of muddy deposits, partly occupied by salt marshes and mangroves, in estuaries, lagoons and the inner parts of Westernport Bay and Corner Inlet. Thus some sectors have receded as the result of erosion, while others have advanced (prograded) by deposition. Coastline evolution has also been influenced by fluctuations in the relative levels of land and sea along the Victo­rian coast in Pleistocene and Holocene times (Jenkin 1981). There is evidence that the sea has been sometimes higher and sometimes lower than it is now. Episodes of higher relative sea level are indicated by emerged beaches and shore platforms above the level at which such features

> The

Western District (Discovery Bay to Mepunga) Port Campbell coast (Mepunga to Princetown) > The Otways and South Bellarine Coast (Princetown to Point Lonsdale) > Port Phillip Bay (Point Lonsdale to Point Nepean) > The Nepean and Flinders Coast (Point Nepean to West Head) > Westernport Bay with French Island and Phillip Island > South Gippsland, Wilsons Promontory and Corner Inlet (San Remo to Corner Inlet) >  East Gippsland (Corner Inlet to Cape Howe) >  The Gippsland Lakes > The

The coastline of Victoria (Bird 1993) is about 1,700 km long, with over 1,000 km of outer coastline, exposed to strong wave action from the Southern Ocean, and more sheltered embayments: Port Phillip Bay (coastline 262 km), Westernport Bay (coastline 263 km) and Corner Inlet (coastline 150 km). It has a variety of coastal landforms: cliffs and bluffs occupy 46% of the coast, sandy beaches 42%, and salt marshes and mangrove swamps 12%. The geological structure of Victoria is the result of the uplift in Tertiary times of the Eastern Highlands of Australia, which run in from New South Wales then swing westward, declining and fading into the hill country around Ballarat and Bendigo, and uplift of the Cretaceous rocks of the Otway Ranges and South Gippsland Highlands. There has been subsidence of intervening areas, notably the Latrobe Valley syncline and the sunklands of Port Phillip, Westernport and Corner Inlet. There have also been contributions from volcanic activity, particularly the outpourings of Older Volcanics in the Eocene and Oligocene, and of the Newer Volcanics, extensive in the Western District, in the Pliocene and Pleistocene (Hills 1975). This extensive Cainozoic vulcanicity is unusual, as western Victoria is not close to the edge of a tectonic plate; it was evidently on a subcrustal ‘hot spot.’

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

⊡⊡ Fig. 21.6.1 Coastal geology of Victoria. (Courtesy Geostudies.)

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21.6 Victoria Introduction

Victoria Introduction

are now forming, Episodes of lower sea level are indicated, by beach and dune formations now out on the sea floor and by the fact that the mouths of river valleys have been submerged by the sea to form inlets, estuaries and lagoons. These changes represent an interaction between tectonic (upward or downward movements of the Earth’s crust) and eustatic movements (rises and falls of sea level, including oscillations resulting from changes in the volume of water in the oceans as the Pleistocene glaciers and ice sheets waxed and waned). The sequence of relative changes in land and sea level can be determined for particular sectors of coastline, but it can be difficult to separate the movements of the land (generally localised) from movements of the sea (regional and global). Earthquakes, the expression of tectonic activity, have been recorded on several parts of the Victorian coast, notably around Corner Inlet, Westernport Bay, Port Phillip Bay and the Otways coast, but the pattern and scale of Holocene uplift and depression are not yet known. Victoria was certainly subject to eustatic oscillations during the Pleistocene, when the world’s oceans rose to high levels during mild (Interglacial) phases and fell to low ­levels during the intervening cold (Glacial) phases. According to Gill (1971) the sea stood about 7.5 m above its present level 120,000 years ago, and about 3 m

21.6

above its present level about 80,000 years ago, both during the mild Last Interglacial climatic phase. It then fell well below its present level during the Last Glacial phase of the Pleistocene, when large quantities of water were withdrawn from the oceans and deposited as ice in spreading glaciers. About 18,000 years ago, when the glaciers and ice sheets reached their maximum extent, the sea was 120– 140 m below its present level, and the continental shelves were exposed as wide coastal plains: the coastline off Victoria at this stage is shown in >Fig.  21.6.2. Then the Earth’s climate became warmer, the glaciers and ice sheets diminished, and a world-wide sea level rise, known as the Late Quaternary marine transgression brought the sea up to about its present level during the Holocene, about 6,000 years ago. The landforms of the Victorian coast have thus been shaped largely during the phase of Holocene sea level stillstand within the past 6,000 years, although there are some features that survive from earlier (Pleistocene) coastlines. It is possible that the sea rose a metre or two above its present level between 4,000 and 6,000 years ago, and then fell back; but landforms and deposits suggestive of this slightly higher Holocene sea level could be the outcome of tectonic uplift of the Victorian land margin rather than a eustatic oscillation. Some of these features may

⊡⊡ Fig. 21.6.2 The coastline of south-eastern Australia 18,000 years ago, when the sea stood 120–140 m below its present level. Victorian rivers then flowed down to the lowered coastline, and a depression on the floor of Bass Strait may then have been a lake basin (Jennings 1959). (Courtesy Geostudies.)

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Victoria Introduction

have been produced by storm surges with the sea at its present level. After the sea arrived at about its present level, parts of the coast have been cut back as cliffs during the Holocene stillstand, while others have been built forward by the accretion mainly of sandy sediment (beaches, spits and barrier formations) or muddy deposition (salt marshes, mangroves, estuarine mudflats and coastal plains). The various coastal landforms have been shaped by processes active in nearshore waters, which include waves produced by storms far away in the Southern Ocean, arriving as ocean swell from the SW and the SE (>Fig. 21.6.3). They are widely spaced, with wave periods (between the passage of  successive wave crests) of 10–16 s, and they produce

­ reakers up to 3 m high. In addition there are locally genb erated waves, produced by onshore winds (mainly southwest to southerly but occasionally south-easterly) in coastal waters. These are typically short, with wave periods of less than 6 s, and during storms they can form breakers several metres high. Variations in wave energy along the Victorian coast are related to the degree of exposure of each sector to ocean swell and storm waves. These can be expressed in terms of significant wave height, defined as the average height of the highest one-third of waves recorded over a stated period. High wave energy coasts can be defined as  those where estimated mean annual significant wave height exceeds 1.0 m, low wave energy coasts where it is

⊡⊡ Fig. 21.6.3 The predominant patterns of ocean swell in Bass Strait are produced by a south-westerly swell entering on either side of King Island and a south-easterly swell (which is the south-westerly swell refracted around Tasmania) moving in towards the Ninety Mile Beach (after Davies 1972). Curving beach outlines have been shaped by refracted ocean swell. East of Wilsons Promontory the sea is notoriously choppy as a result of intersecting wave patterns. (Courtesy Geostudies.)

Victoria Introduction

less than 0.3 m, with moderate wave energy coasts being the intermediate category. Examples of high wave energy coast in Victoria are found on the southern coast of the Portland Peninsula, where a mean annual wave height of  2.7 m has been recorded off Point Danger, between Warrnambool and Cape Otway, and at Cape Schanck, Cape Liptrap and Wilsons Promontory. The eastern and western coasts of Port Phillip Bay are moderate and low wave energy coasts respectively. The range of tides in Victorian coastal waters is generally less than 2 m during fortnightly maximum (spring) tides in Bass Strait. Tide range increases into inlets and embayments where the tidal flow is magnified by interacting with narrowing configuration, attaining more than 3.3 m towards the head of Westernport Bay and 2.7 m at the mouths of rivers draining to the northern shores of Corner Inlet. The tide range in the northern part of Bass Strait is 1.6 m, but because of the narrow entrance at Port Phillip Heads, spring tides diminish to 1.1 m at Point Lonsdale and about 0.6 m at Williamstown, at the head of Port Phillip Bay. The intertidal zone is typically 50–100 m wide along the ocean coast and 20–30 m wide around Port Phillip Bay. In Westernport Bay, the area of marshes, mudflats and sandflats exposed at low spring tides is about 270 sq km (nearly 40% of the area submerged at high spring tides). Corner Inlet and the region behind the barrier islands to the east have an intertidal zone of about 180 sq km, Port Phillip Bay about 28 sq km and Andersons Inlet, the largest estuary in Victoria, about sq 16 km. The total intertidal area in Victoria is about 600 sq km.

21.6

Beaches are extensive on the coast of Victoria, and are generally gently curved in outline, shaped by refracted swell on the outer coast and locally generated waves in Port Phillip Bay and Westernport Bay. Most beaches are sandy, but some are shelly or gravelly. Parts of the Victorian coast have been modified by the building of sea walls and groynes intended to halt erosion and retain beaches, and by the construction of harbour breakwaters and other shore structures. Coastal land reclamation has proceeded locally, mainly in port areas, and some cliff sectors have been artificially stabilised. Engineering structures have in some places caused erosion and accretion that would not otherwise have occurred. Human activities have also modified the vegetation and topography of coastal dunes, salt marshes, mangroves and seagrass areas, in ways that will be discussed in subsequent chapters.

References Bird ECF (1993) The coast of Victoria. Melbourne University Press, Melbourne Davies JL (1972) Geographical variation in coastal development, Oliver and Boyd, Edinburgh (2nd edn, Longman, London, 1980) Gill ED (1971) The far-reaching effects of Quaternary sea level changes on the flat continent of Australia. Proc R Soc Vic 84:189–205 Hills ES (1975) Physiography of Victoria, 2nd edn. Whitcombe and Tombs, Melbourne Jenkin JJ (1981) Evolution of the Victorian coastline. Proc R Soc Vic 92:37–54 Jennings JN (1959) The submarine topography of Bass Strait. Proc R Soc Vic 71:49–72

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21.6.1 Victoria: The Western District (Discovery Bay to Mepunga)

1.  Introduction The Western District of Victoria is dominated by the wide lava plains of the Plio-Quaternary Newer Volcanics, which outcrop in coastal headlands on the Portland Peninsula and around Port Fairy. Along the coast these are bordered and overlain by Pleistocene and Holocene sand deposits in successive embayments backed by beaches and dunes.

2.  The Coastline The gently curving coast of Discovery Bay, 60 km long, faces SW, and has a sandy surf beach backed by extensive dunes, some fixed beneath grassy or scrub vegetation, others bare and mobile. The beach is dominated by creamyyellow fine to medium calcareous sand, with 60–80% calcium carbonate content, the balance being mainly quartz: there are also patchy splays of shelly gravel, and a few pebbles of sandstone and flint. The dunes are of similar sand, blown landward from the beach. Some of the older (Pleistocene) dunes have been lithified by secondary precipitation of carbonates in cavities between the sand grains, and so converted into the dune sandstone or sandy limestone known as dune calcarenite. The coastal dune formations occur at the convergence of dune calcarenite ridges extending from South Australia (>Fig. 21.6.1.1) into an area of transgressive dunes. The Pleistocene dune calcarenites behind Discovery Bay are arranged in overlapping sequences, the Lower Pleistocene dunes forming a truncated basement against which Upper Pleistocene dunes, also lithified as dune calcarenites, accumulated before the younger, unconsolidated Holocene dunes were banked up along their seaward fringe (>Fig. 21.6.1.2). In the hinterland the Pleistocene dunes mantle a plateau into which Glenelg River has incised a deep gorge, with cliffs up to 25 m high in stratified Miocene limestone. It runs west, then south to Nelson and the sea. The Holocene dunes form a seaward fringe behind the sandy beach, with vegetated foredunes up to 15 m high in the

NW, interrupted by blowouts spilling landward. The beach interrupted is by a few outcrops of the underlying Pleistocene dune calcarenite as low cliffs and short segments of shore platform at McEacherns Rocks, Noble Rocks, and Sutton Rocks (Cape Montesquieu). To landward are extensive mobile dunes, which were held by scrub and woodland vegetation until this was destroyed by grazing and burning in the nineteenth century. Amid the dunes are hollows containing dune lakes. Johnstone Creek descends an escarpment and flows through a chain of depressions which contain lakes of ­varying depth and size, much influenced by preceding rainfall (>Fig. 21.6.1.3). The most permanent of these, Swan Lake, has an extensive fringe of freshwater sedges and scrub. Further east the scrub-covered dunes behind the ocean beach are backed by Bridgewater Lakes (>Fig. 21.6.1.4). These were formerly brackish lagoons, but they have been completely cut off from the sea by the accretion of wind-blown sand, and are now fresh, with reeds and rushes growing around their shores. They are backed by a steep slope of Pleistocene dune calcarenite banked over Plio-Pleistocene basalts. It contains the large Bridgewater Caves, with well-developed calcareous dripstone structures beneath a capping of Pleistocene dune calcarenite. To the south the steep slope curves laterally into a coastal cliff, and as it does so the dune calcarenite is undercut to expose basal outcrops of volcanic ash (tuff) and dark green basaltic lava on the shores of Descartes Bay. The Portland Peninsula is an upland on which undulating dunes and dune calcarenite ridges overlie a plateau composed of Newer Volcanics, the internal structure of which is exposed in the coastal cliffs. Behind the cliff crest is a so-called petrified forest, where tree trunks and root structures preserved in calcrete have been exposed by the erosion of the surrounding dune sand. Cape Duquesne is the first of a series of headlands, cut into an array of volcanic rocks and structures, including pahoehoe (ropy) and columnar lava. On the west coast of Bridgewater Bay the cliff has been cut across a massive Pleistocene volcano. Boutakoff (1963) suggested that the

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

21 . 6 . 1

Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.1 Pattern of uplifted Pleistocene dune calcarenite ridges in the SE of South Australia, and their convergence behind Discovery Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.1.2 The relationship between successively deposited Pleistocene dune calcarenite formations and capping Holocene dunes in front of the fault-line scarp in the coastal region of Discovery Bay (after Boutakoff 1963). (Courtesy Geostudies.) SCARP LAKE

BEACH Holocene dunes PRESENT SEA LEVEL

T

Pleistocene dune calcarenites

Pleistocene dune calcarenites

Basalt

FAUL

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Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.3 Mobile dunes advancing inland from the shores of Discovery Bay. Johnstone Creek drains into Swan Lake, which has an overflow channel through the dunes during wet periods. (Courtesy Geostudies.)

bays (Bridgewater Bay, Nelson Bay and Grant Bay) were essentially large volcanic craters (calderas) that had been invaded and submerged by the sea. Nelson Bay retains sediment that were deposited in a freshwater crater lake which must have been separated from the sea by a volcanic rim that has since been breached; they are now cut back as cliffs by the sea. There is a shelly beach in Bridgewater Bay (Gell 1978), backed by scrub-covered dunes that were formerly driven towards Portland by the prevailing SW winds. The cliffs at Cape Nelson expose more volcanic plugs and stratified lavas and tuffs, with caves cut out along joints and bedding planes. Cape Sir William Grant also has cliffs cut in volcanic structures, modified on the eastern side by

quarrying. At several points around the Portland Peninsula there are features which were shaped when the sea stood at a higher level, relative to the land. There are rock benches up to 5 m above present sea level on the Capes, and an emerged shore platform at Point Danger (>Fig. 21.6.1.5), which has gravely deposits carrying a sparse vegetation, and is occasionally overwashed by storm waves at high spring tides, but is essentially a relict feature. Some of the features formed during the Late Pleistocene high sea level phase and were revived when the sea returned to a similar level in the early Holocene (Bird 2008). Offshore, Lawrence Rocks consist of three rugged high islands of volcanic rock, the last remains of the eastern rim of a largely submerged caldera beneath Nelson Bay. They

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Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.4 Mobile dunes advancing inland from the shores of Discovery Bay. Johnstone Creek drains into Swan Lake (S) in the foreground. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.1.5 The emerged beach at Point Danger, south of Portland, backed by a bluff which was a receding cliff when the sea stood at a slightly higher level earlier in Holocene times. (Courtesy Geostudies.)

include a flat-topped ridge of bedded tuff draped with white guano dropped by the many sea birds that dwell on these rocks. The east coast of the Portland Peninsula is relatively sheltered, with bluffs that decline northward to Battery Point, where they are fronted by an artificial boulderfringed reclaimed area extending to the Portland harbour breakwater. This was built between 1957 and 1961 by dumping basalt blocks quarried from Cape Sir William Grant. It

resulted in local sand accretion, some of which was used to renourish Nunns Beach, to the north (>Fig. 21.6.1.6). To the north are the limestone cliffs at Whalers Bluff, up to 10 m high, cut into white chalky Portland Limestone, of Miocene age (> Fig. 21.6.1.6). Overlying Pliocene clays and marls, sandy limestones and weathered Pleistocene basalt form a gentler slope, largely scrub-covered. The basalt has weathered to form a sticky brown clay which cracks when dry and expands when it is soaked. It is

21 . 6 .  1

Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.6 Nunns Beach, Portland, after it was replenished with fine sand in 1990. Winds then piled the fine sand into low hummocky dunes which were colonised by grasses and shrubs. In the background the slope-over-wall profile at Whalers Bluff consists of soft PlioPleistocene formations over white Portland Limestone, the cliff base being protected by a boulder wall. (Courtesy Geostudies.)

unstable, ­particularly along the crest of the cliffs, and there has been recurrent slumping of the coastal slope. In 1975, a major landslide occurred, when a large mass of weathered basaltic, clay and underlying calcareous rubble collapsed into the sea. The cliffs of the Portland coast become grassy bluffs where the coast road (Dutton Way) descends to the shore. They then curve behind the coastal plain, are interrupted by the valley of the Surry River, can be followed until they fade out on the western side of the Fitzroy valley, and revive in the country behind Port Fairy to eventually intersect the present coastline at Warrnambool. These bluffs were formerly marine cliffs, a legacy of a Late Pleistocene phase when the sea stood at a slightly higher level about 80,000–100,000 years ago. In recent decades the erosion that has become a problem along the sandy shore below Dutton Way may have resulted from modification of the local wave regime and the interception of northward drifting sand by the building of the Portland Harbour breakwater. The response to beach erosion has been to armour the shore with blocks of basalt and dune calcarenite, but this has led to further lowering and depletion of the sandy beach and nearshore area by scour caused by wave reflection from the boulder rampart. The Surry River has incised a channel through a dune calcarenite ridge at Narrawong, and then flows along a swampy swale before curving round to enter the sea. The valley of Fitzroy River was modified by lava flows during the last phase of eruption of Mount Eccles, some 25 km

inland. Lava then flowed southward to the coast, then out over what is now the sea floor for a distance of 16 kim, terminating 37 m below present sea level (Boutakoff 1963). Between Codrington and Yambuk the sandy beach is backed by a grassy Holocene foredune, backed by a swampy corridor (formerly an elongated coastal lagoon), then a series of roughly parallel ridges of Pleistocene dune calcarenite. Eumeralla River flows eastward until it enters Lake Yambuk, an estuarine lagoon which also receives the Shaw River. In the hinterland the underlying Newer Volcanics emerge to form hummocky basaltic country, the Fingerboards. Extensive reed and rush swamps fringe Lake Yambuk, a lagoon which has a sandy outlet channel winding to the  sea between grassy and scrubby dune ridges (>Fig. 21.6.1.7). To the east the coastal dune fringe becomes higher, the underlying Pleistocene dune calcarenites emerge to form low rocky cliffs and shore platforms at The Crags. This substantial sector of dune calcarenite coast shows headlands of thinly bedded dune brown sandstone capped by harder grey calcretes and interbedded ancient soils (palaeosols). The cliff crest has a stripped edge, exposing knobbly calcrete from which overlying dune sand has been removed by wind scour. Between the headlands sandy coves have been cut out in dune hollows. Prominent offshore is Lady Julia Percy Island, an almost flat-topped slab of Newer Basalt, fringed by steep, often vertical, cliffs. To the east low, rugged calcarenite cliffs continue to Cape Reamur, where basalt of the underlying Pleistocene Newer Volcanics

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21 . 6 . 1

Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.7 The outflow channel from Lake Yambuk curving westward behind a dune-capped barrier spit. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.1.8 Erosion on the shores of Belfast Lough, a coastal lagoon near Port Fairy. (Courtesy Geostudies.)

appears on a boulder-strewn shore. Goose Lagoon is a swampy area in a former marine inlet. Behind Cape Reamur the Newer Volcanics form an extensive gently undulating plain, with knobs of lava and many small lakes and rushy hollows. East of Cape Reamur it outcrops on the shore as a black, boulder-strewn platform, sloping gently seaward and subdivided by well-defined horizontal and vertical joints into columnar structures of varying size. At Port Fairy the River Moyne has outlets west and north of Rabbit Island, a rocky island with dunes capping

a dissected lava flow. It has passed through Belfast Lough, a shallow lagoon behind the high Holocene dunes that back Port Fairy Bay. Its former reedswamp fringe has disappeared following increased salinity in the lagoon, which has led to its replacement by salt marsh. Poles mark the remains of a former wooden shore wall (>Fig. 21.6.1.8). Port Fairy Bay has a long sandy beach, interrupted by spurs of disintegrated lava and backed by grassy dunes. It had here that marram grass was first planted (in 1883) to stabilise coastal dunes in Victoria. The coastal plain is

21 . 6 .  1

Victoria: The Western District (Discovery Bay to Mepunga)

⊡⊡ Fig. 21.6.1.9 Layers of volcanic ash (tuff) from the Tower Hill volcano, exposed in a cliff on the coast west of Warrnambool. (Courtesy Geostudies.)

backed by an escarpment cut into the Miocene Port Campbell Limestone: an Early Pleistocene cliffed coast. To the east the escarpment is breached by the large volcanic crater (a caldera) at Tower Hill, the inner wall of which exposes thin layers of tuff dipping away at low angles from the former eruption centre. These dominate the surrounding country, and extend to the coast near Warrnambool, to the east. To the south, Armstrong Bay has a sandy beach, backed by a wide ridge of high grassy dunes which are thought to conceal the remains of a sixteenth century ‘Mahogany Ship’, a Portuguese caravel (>Fig. 21.6.1.9). The Holocene dunes continue SE, underlain by Late Pleistocene dune calcarenites which eventually emerge as low cliffs and shore platforms, with intervening sandy coves towards Warrnambool. Behind the dune fringe is a broad swampy corridor through which flows the Merri River, deflected SE into Lady Bay (Gill 1984). A stone breakwater completed in 1890 shelters Warrnambool Harbour here, and subsequently sand accretion has formed a wide sandy beach backed by low grassy dunes. These are backed by freshwater Lake Pertobe amid marshes in a hollow in the dune calcarenite fronting the Warrnambool bluff. Lady Bay curves round to the mouth of Hopkins River, which comes down through a gorge incised into the Port

Campbell Limestone, with outcrops of horizontally bedded strata in the bordering river cliffs. South-east from the mouth of Hopkins River is Logan Beach, backed by dunes 10–15 m high, with a steep, generally vegetated seaward slope, on which a whale-watching platform stands. Along the coast the dunes are cliffed, then pass into steep and rugged cliffs cut in a high coastal ridge of Pleistocene dune calcarenite, with deep clefts and caves, but only intermittent shore platforms. South of Mepunga the Port Campbell Limestone begins to outcrop as windows in the rugged dune calcarenite cliffs, and eventually emerges to form vertical cliffs, with a capping of dune calcarenite persisting towards Childers Cove. It is here that the magnificent cliff scenery of the Port Campbell coast begins.

References Bird ECF (2008) Coastal geomorphology, 2nd edn. Wiley, Chichester Boutakoff N (1963) The geology and geomorphology of the Portland area. Geol Surv Victoria. Memoir 22 Gell R (1978) Shelly beaches on the Victorian coast. Proc R Soc Vic 90:257–270 Gill ED (1984) Coastal processes and the sanding of Warrnambool Harbour, Warrnambool Institute Press, Warrnambool

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21.6.2 Victoria: The Port Campbell Coast (Mepunga to Princetown)

Introduction The 50-km coastline between Childers Cove and Glenample consists of steep, often vertical and sometimes overhanging cliffs of up to 70 m high, cut into soft Port Campbell Limestone. It is exposed to high wave energy from the prevailing SW ocean swell and storm waves arriving through deep water over a narrow section of the Australian continental shelf, here only about 60 km wide. The sea is very rarely calm. Even on windless days, a long, even ocean swell generated out in the Southern Ocean rolls in, waves breaking on the shore every 10–16 s. In a SW gale huge storm waves crash against the cliffs, sometimes surging up more than 30 m to their crests. Mean spring tide range is small (about 1.2 m at Port Campbell), so that the large waves consistently attack the base of the cliffs. The Port Campbell coast can be treated as a unit developed on a single geological formation (Baker 1943), the Port Campbell Limestone, of Miocene age, consisting of stratified yellow-brown calcareous clays, silts and thin sandy limestones (>Fig.  21.6.2.1). East of Mepunga the dune calcarenite capping fades out along the cliffs, which are backed by a gently undulating treeless healthy plateau that in places rises more than 30 m above sea level. Generally the cliff crest is even, except where it declines towards the incised river valleys, but in the vicinity of Childers Cove, where the ground slopes inland, there are high headlands and intervening coves with lower bayhead cliffs. Because of the landward decline from the cliff crests, these cliffs are diminishing in altitude as they retreat. Inland, the wide Nullawarre plateau is dotted with numerous sinkholes (swampy depressions). Near Stanhope Bay the cliff crest shows a calcrete horizon exposed where overlying dune sands have been removed, and solution processes have weathered it into rugged, knobbly, karstic (produced by solution processes) topography. Behind Three Mile Beach a formerly vertical cliff has been cut off from marine erosion by the accumulation of beach and dune sand. It has been modified by subaerial degradation, and shows grassy bevels and ­convexities on the crest and concave basal fans of

downwashed sediment, with an intervening residual cliff face. If these processes continue the former vertical cliff will be re-shaped into a convex-above-concave slope profile similar to that of valley sides inland (Bird 1977). In the Bay of Islands marine dissection of the receding cliffs has isolated numerous scattered stacks, which persist ⊡⊡ Fig. 21.6.2.1 The cliff in Port Campbell Limestone at Goudies Lookout has a series of ledges on the outcrops of hard horizontal strata. The scale is given by the man (arrowed) on the cliff-top, who has a fishing line down into the surf. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.6.2 

Victoria: The Port Campbell Coast (Mepunga to Princetown)

offshore in various stages of reduction to flat platforms, usually just below low tide level. At Peterborough the cliffs are interrupted by the mouth of Curdies Inlet, a shallow lagoon bordered seaward by the dune-capped sand barrier behind Newfield Bay. The lagoon is fringed by reeds and rushes, which form an extensive swamp at the northern end, where the Curdies River flows in from a wide valley incised into the Port Campbell Limestone plateau. The outlet from Curdies Inlet is often sealed off by the accretion of beach sand, but after rainy weather the lagoon level rises, and overspill cuts a new outflow channel. In some years this process has been hastened by digging a trench through the sand barrier to initiate outflow. When the outlet is open the lagoon level falls, and a sandy threshold bank is exposed. The Grotto occurs where one of the sinkholes on the coastal plateau has been intersected by cliff recession to form a hole floored with fallen boulders and opening seaward through an archway to a ledge with a marine pool, a few metres above sea level. The pattern of cliff recession is much influenced by vertical joining, mainly NW–SE and NE–SW. The cliff face shows breakaways along joint planes, and in the bay near London Bridge high rectilinear columns of horizontally bedded strata have broken away from the cliffs, leaving a serrated cliff edge (>Fig. 21.6.2.2). London Bridge used to be a fine example of a double natural archway formed on an elongated promontory where the sea had cut out two caves between bordering joint planes (>Fig. 21.6.2.3). On 15 January 1990 the inner arch collapsed into the sea, leaving a heap of blocks and boulders that has subsequently diminished. The outer arch persists, standing offshore. Much of the coastal outcrop of Port Campbell Limes­ tone is capped by a red-brown clay varying in thickness of up to 6 m, and containing small iron concretions known as buckshot gravel. The clay outcrop sometimes forms a bevel above the vertical limestone cliffs, as at Point Hesse. This post-Miocene deposit was described as the Hesse Clay by Gill (1976). Its contact with the underlying Port Campbell Limestone is often uneven, lowered into depressions that have formed where percolating groundwater has dissolved the limestone. Some of the smaller streams on the coastal plateau disappear down sinkholes near the edge of this impervious mantle of clay. The Arch east of Point Hesse (also known as Marble Arch) is a ­natural arch on a rock ledge 6 m above the sea. It is still much as it was when Baker (1943) photographed it 65 years ago, the rock surface having been indurated by carbonate precipitation. The coast at Two Mile Bay (>Fig. 21.6.2.4) is of great interest because it preserves a fragment of the Late Pleistocene landscape and throws light on the evolution of

⊡⊡ Fig. 21.6.2.2 Subsiding columns of joint-bounded rock on the cliff near London Bridge. (Courtesy Geostudies.)

the coasts bordering Bass Strait. During the Last Glacial ­low-sea level phase bluffs of this kind must have extended along the whole of the formerly cliffed coastline. Aborigines who came here about 40,000 years ago would have seen these bluffs as an inland slope, descending to the wide plains that led them across to Tasmania: a landscape probably largely covered by forest, scrub and heath. The bluffs have not yet been rejuvenated by marine erosion, which has cut cliffs along the coastline to the east and west. The bluff behind Two Mile Bay declines to a narrow swampy lowland terrace bordered by a dune ridge on its seaward side, capping low cliffs cut in Port Campbell Limestone. There is a visor and notch overlooking ­segments of beach, ramps of abraded limestone, and a modern subhorizontal shore platform with a veneer of beach sand. The dunes are underlain by an emerged shore platform, and the bluff to landward is a former sea cliff, degraded by subaerial processes that have formed a colluvial apron (>Fig. 21.6.2.5). Its profile is similar to the slopes inland on the sides of the incised Port Campbell valley.

Victoria: The Port Campbell Coast (Mepunga to Princetown)

21.6.2

⊡⊡ Fig. 21.6.2.3 London Bridge as it was in 1989, before the inner arch collapsed in 1990. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.2.4 An aerial view of Two Mile Bay, Port Campbell, showing the Pleistocene bluff running behind the dune-fringed coastal terrace. (Courtesy Neville Rosengren.)

East and west of Two Mile Bay the bluff has been rejuvenated as receding cliffs, and its preservation in Two Mile Bay is because of the presence of a nearshore limestone reef, which diminishes wave attack on this sector. Because of this reef, swell breaks into surf farther offshore than on other parts of the coast. The reef is thus a protective feature which has preserved a segment of Late Pleistocene coastline that would have to be cut back about 300 m to convert the bluffs into vertical cliffs. In general shore platforms are poorly developed on the Port Campbell coast, and where they do occur they are

usually structural, in the sense of harder limestone payers exposed by marine erosion in front of a receding cliff. Port Campbell is a valley-mouth inlet, bordered by cliffs that decline landward into bluffs on either side of a small crescentic sandy bay-head beach. The cliffs are penetrated by caves at the mouths of underground tunnels. There was a major landslide east of Port Campbell in 1939 when a section of cliff near Sentinel Rock 70 m long and up to 12 m wide suddenly fell into the sea. The scar of this fall can still be seen, and after 70 years only part of the tumbled rock has been consumed by marine erosion.

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21.6.2 

Victoria: The Port Campbell Coast (Mepunga to Princetown)

⊡⊡ Fig. 21.6.2.5 Cross-section across the coastal terrace at Two Mile Bay, Port Campbell, based on diagrams by Baker and Gill (1957). (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.2.6 A cliff collapse near The Amphitheatre in 1970. (Courtesy Geostudies.)

A fall of this magnitude probably occurs on each sector of these cliffs once in 1,000 years, but it can happen at any time: there have been many smaller rock falls. The present cliffed coast has a generally sinuous or crenulate outline, with numerous small bays and narrow gorges. Goudie’s Lookout (>Fig.  21.6.2.1) is a narrow ­promontory 128 m long, 6–12 m wide and up to 60 m high, with prominent layers of more resistant limestone etched in the cliff. Most cliff falls go unrecorded, but the cliff collapsed at The Amphitheatre in 1970 (>Fig. 21.6.2.6) and Elephant Rock, the long, narrow promontory behind Mutton Bird Island, lost its “trunk” during a storm in 1935 (Baker 1943). In many places the crests of the vertical cliffs have been stripped bare of soil and vegetation by wind erosion and

the effects of waves and spray during storms, as well as runoff during heavy rain (Baker 1958). The bald-topped Baker’s Oven Rock had soil and weathered material swept away by wave overwash, and stripped zones are prominent on cliffs either side of Sherbrook River. Photographs published by Baker (1943) showed storm waves breaking over a 25-m high headland at Broken Head. The backwash following such giant waves sweeps loose material away from the cliff top and cliff face, and as a result there is a cliff crest ledge up to 50 m wide, backed by a low concave bluff of brown clay up to 4 m high at the trimmed edge. By contrast, some slightly more sheltered sectors of the cliffed coast have a cliff-top deposit of yellow-brown calcareous marl (parna) as a wedge or levee up to 0.5 m thick and 100–150 m wide, thinning and fining landward. Ashton et al. (2002) concluded that this marl was derived from the

Victoria: The Port Campbell Coast (Mepunga to Princetown)

cliff face by wind action, and carried up and over the cliff crest. It contains Miocene fossils derived from the cliff outcrops. The marl has been deposited on trees in bark furrows and branch angles. It rests upon a buried podzolic soil that emerges landward, and carries ­calcicolous vegetation that passes into calcifugous heath on the podzolic soils. East of Broken Head the cliff outlines consist of alternations of deep narrow gorges, such as Survey Gorge, and long narrow promontories, both bordered by vertical cliffs. Island Archway is an elongated stack penetrated by a wide arch, the whole formation strongly influenced by erosion along joint planes. The arch collapsed into the sea in June 2009. Mutton Bird Island is another block penetrated by a natural arch and the Razorback is a long, high and remarkably narrow ridge to the east of Lochard Gorge. This is another steep-sided inlet between rectilinear promontories, with a beach of in-washed sand (>Fig. 21.6.2.7). East of The Razorback the vertical cliffed coastline straightens, and is bordered by a group of tall stacks known as The Twelve Apostles, with outlines related to patterns of jointing and stratification (>Fig.  21.6.2.8). In the lee of these are minor protrusions, and it is possible that several of these were formerly linked to the adjacent stacks by arches that have since collapsed. One of the stacks collapsed in 2006. Beaches are poorly developed on the Port Campbell coastline, due partly to a meagre sand supply and partly to wave reflection from the vertical cliffs, which prevent beach accretion. The rivers that drain to this coast carry little sediment into the sea, and the material derived from cliff erosion is generally too soft and too fine in texture to ⊡⊡ Fig. 21.6.2.7 Lochard Gorge. (Courtesy Geostudies.)

21.6.2

be retained on beaches exposed to strong wave energy: it is silty, soon dispersed and carried well out to sea. Beaches do occur where there is, or has been, sandy dune ­calcarenite on the cliff crest, as in the Childers Cove area and near Gibson Steps. As the cliffs recede, quartzose and calcareous sand from this capping material falls to the shore, and has been incorporated in the beach. The general absence of beaches has enabled storm waves to attack the base of the cliffs vigorously and consistently. Cliff recession would have been slower had wider, more protective beaches been able to accumulate along this coastline. Pleistocene dune calcarenite appears east of Gibson Steps, where it is banked against an eroded bluff that backs a scrubby landslide where Port Campbell ­Lime­stone has collapsed behind a re-entrant on Gellibrand Clay (>Fig. 21.6.2.9). The dune calcarenite forms a high and wide coastal ridge on the southern side of the Latrobe Creek valley. It descends gradually SE to the shore to outcrop in the cliffs towards Point Ronald, and is truncated by a steep cliff 90 m high, alongside the mouth of Gellibrand River. It is possible to distinguish six successive dune formations, separated by wavy unconformities that represent episodes when dune accretion was interrupted by wind erosion, truncating the earlier dune bedding. Three of the un-­ conformities include palaeosol horizons and associated underlying calcrete layers that mark stages when the land surface became sufficiently stable for soils to form. Gellibrand River forms a convenient boundary between the Port Campbell coastline and the rising slopes of the Otway Ranges.

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21.6.2 

Victoria: The Port Campbell Coast (Mepunga to Princetown)

⊡⊡ Fig. 21.6.2.8 The 12 Apostles. (Courtesy Jenifer Bird.)

⊡⊡ Fig. 21.6.2.9 The Glenample Landslide, east of Gibson Steps. (Courtesy Geostudies.)

References Ashton DH, Williams RJ, McDonald M (2002) The ecology of cliff-top heathlands at Port Campbell, Victoria. Proc R Soc Vic 114:22–41 Baker G (1943) Features of a Victorian limestone coastline. J Geol 51:359–386 Baker G (1958) Stripped zones at cliff edges along a high wave energy coast, Port Campbell, Victoria.  Proc R Soc Vic 70:175–179

Baker G, Gill ED (1957) Pleistocene emerged platform, Port Campbell, Victoria. Quaternaria 4:55–68 Bird ECF (1977) Cliffs and bluffs on the Victorian coast. Vic Nat 94:4–9 Gill ED (1976) Warrnambool-Port Fairy district. In: Douglas JG, Ferguson JA (eds) The geology of Victoria, Vol 5. Geological Society of Victoria, Special Publication, Australia, Melbourne, pp 299–304

21.6.3 V  ictoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

1.  Introduction The Otways coast extends from the mouth of Gellibrand River at Princetown SE to Cape Otway, then NE past Lorne to Anglesea and Barwon Heads. It truncates the Otway Ranges, which consist of uplifted Cretaceous sedimentary formations flanked by the overlying Tertiary rocks, which dip away on either side and also occupy an interior trough. Consequently, the coastline may be divided on geological grounds into five segments: the Tertiary outcrops between Princetown and Moonlight Head, the Cretaceous rocks to Johanna River, the Tertiary trough between Johanna and Aire River, the Cretaceous rocks again from there past Cape Otway to Eastern View, and finally the Tertiary outcrops from Eastern View to Barwon Heads. The country is higher and more deeply dissected by river valleys than the coastal plateau of the Port Campbell district.

2.  The Coastline South-east from the mouth of Gellibrand River the dunebacked sandy beach gives place to cliffs with small promontories and bays in front of an ascending coastal ridge, backed by Rivernook and the Gellibrand River valley. The ridge consists of Miocene sands, clays and gravels with a capping of Pleistocene dune calcarenite and Holocene dunes, but there are few outcrops on the steep scrubby slopes until the cliffs begin at Buckleys Point, about a kilometre SE of the mouth of Gellibrand River. These cliffs show sections in gently westward tilted Tertiary strata as they begin to rise on the flank of the Otway Ranges. The Gellibrand Marl (Miocene) is underlain by a series of younger formations that rise successively to outcrop along the coast SE of Buckleys Point. Oligocene gritty sandstone and limestone are underlain by clays and marls at Buckleys Point, beneath which thick Eocene carbonaceous sandy clays appear in the cliffs. Under this are Palaeocene ferruginous quartzose grits and sandstones

with some pebbly conglomerate. There are patchy nearshore reefs in front of sandy beaches, but no continuous shore platform on these Tertiary outcrops. South-east from Pebble Point the weathered ferruginous conglomerates are prominent. The underlying Cretaceous sandstones and mudstones form shore platforms, then rise in the cliffs which increase in height along the coast towards Bell Point. An upper cliff up to 20 m high develops on the coastal slope up to 400 m inland behind Buckleys Point, and becomes more prominent south-eastward, above Pebble Point and Dilwyn Bay. It continues behind a succession of broad amphitheatres along the coast above Marie Gabrielle Beach as far as The Gable. In each of these the Lower Tertiary formations have slumped seaward, forming an irregular thickly vegetated, hummocky landslide topography. Marie Gabrielle Beach consists of coarse yellow quartz and shell sand is backed by steep bluffs and cliffs cut in grey Cretaceous sandstones and mudstones. A shore platform and vertical cliff cut in Cretaceous sandstones run out to Point Lucton, and the overlying Tertiary formations rest unconformably upon Cretaceous sandstones and mudstones high in the cliffs towards The Gable. The Tertiary sediment disappear on Moonlight Head, 90 m above the sea, and the coastal slopes then consist entirely of Cretaceous rocks. The Cretaceous sandstones and mudstones (also termed siltstones) that outcrop on the coast at Moonlight Head rise to form the high dissected upland ridges of the Otway Ranges (Baker 1950). They have been described as arkoses, a term defined as sandstones with more than 25% feldspar, but they are now known as feldspathic arenites. These Cretaceous formations have been uplifted, folded and faulted, in two broad elongated domes trending SW–NE, one truncated by the coast between Moonlight Head and Johanna, the other from the mouth of Aire River to Cape Otway. Shore platforms are extensive on the Cretaceous rocks of the Otway coast (Hills 1971). At low tide they are exposed for several hours, drying out in calm

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.6.3

  Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

weather, but overwashed by occasional waves when the sea is rough. They are higher and wider on outcrops of massive sandstone, while the softer mudstones have been etched out more rapidly to form intervening bays, often containing sandy beaches. The shore platforms are best developed as horizontal benches on mudstones weathered down to about mean high tide level, and on flat-lying sandstones; the thicker and harder sandstones outcrop as benches or as cuestas (scarps and dip-slopes) on upstanding rocky outcrops, with intervening corridors of planeddown mudstone. In places, there are flat-topped residual mesas of sandstone above the general level of a mudstone platform. Often the shore platforms show a strong structural influence, sloping landward or seaward with the local dip of the rock strata. Disintegration by wetting and drying eventually reduces the rock surface to the level of permanent saturation, where flat platforms are produced: the loosened material is washed away until the surface corresponds with the top of the permanently wet rock. The beaches that occur at intervals along the Otways coast are mainly of yellow-brown calcareous sand with small amounts of clear quartz washed in from the sea floor, but there are also varying proportions of grey felspathic sand and gravel derived from erosion of the cliffs and shore platforms. Much of the material generated in the course of cliff recession is silt and clay produced by the weathering of mudstones, too fine to be retained in the beaches, but locally there are horizons of angular quartz sand and gravel which have disintegrated to contribute to the beach sediment. Streams draining the Otway Ranges have also contributed small quantities of felspathic sand and gravel during occasional flood discharge, but on the sandy beaches the proportion of sediment derived from cliff and foreshore outcrops of Cretaceous rocks, or from fluvial yields, rarely attains 10%. Locally dunes of calcareous sand have formed behind the beaches, and there are segments of dune calcarenite that developed during Pleistocene times. Between Moonlight Head and Cape Otway is a high energy coast, exposed to ocean swell and SW storm waves arriving through deep water close inshore. At high tide these waves break heavily over the shore platforms and attack the cliffs at the base of the steep coastal slopes, which are receding by way of undercutting and recurrent slumping. The bold coast has cliffs up to 90 m high on headlands such as The Gable, Moonlight Head and Lion Headland, and the intervening slopes are very steep, thickly vegetated, and up to 150 m high, trenched by deep valleys such as Walls Gully. There are intermittent shore platforms and some sandy and gravelly coves. The steep coast continues past Maudes Point, with several small

valley-mouth coves, one of which contains sandy Sutherlands Beach, and then declines towards Johanna, where the sandy shore is backed by Holocene dunes. On the western side of the Johanna River valley the Cretaceous sandstones subside as a monocline brings down the Lower Tertiary rock formations which outcrop in the cliffs towards Rotten Point. Here the folded dark grey Cretaceous sandstones are overlain by the lightcoloured Rotten Point Sand (Palaeocene), which is succeeded upwards on the scrubby slopes by the dark red Johanna River Sand and the Browns Creek Clay (both Eocene), and the Castle Cove Limestone (Lower Oligocene). Johanna River valley has been excavated in these soft sedimentary formations, and opens to a beachfringed coast backed by Holocene dunes, with lobes of Pleistocene dune calcarenite running inland. The dune calcarenite outcrops in a low cliff, a spiky stack and a shore platform in the middle of the beach. Near Castle Cove the Eocene sands and clays are almost vertical, but the dip declines quickly SE. The cliffs are then cut in a sequence of Upper Eopcene, Oligocene and Miocene sands, clays and limestones. Pleistocene dune calcarenite, then caps a coastal ridge behind the long sandy beach south-east past the mouth of Aire River, which reaches the coast through a wide gorge incised through Pleistocene dune calcarenite (>Fig. 21.6.3.1). At Point Flinders the Cretaceous sandstones and mudstones reappear and rise towards the Cape Otway anticline. The dune calcarenite that caps the cliffs between Point Flinders and Cape Otway is part of an extensive area of parabolic and formerly transgressive dunes, now stabilised beneath a cover of grasses and scrub, which extends ­eastward across the promontory towards Point Franklin (>Fig. 21.6.3.2). The rising cliffs and undulating shore platforms extend SE and round the bold sandstone promontory of Cape Otway. At Cape Otway the cliffs cut in Cretaceous sandstone are up to 80 m high, with scrubby vegetation on an upper, convex slope of dune calcarenite. The shore platform undulates strongly, with ledges of hard sandstone overwashed by occasional storm waves. The Pleistocene dune calcarenites on Cape Otway consist of aeolian sandstone with calcrete horizons and red palaeosols, overlain by unconsolidated Holocene dunes, mostly stabilised by vegetation, with thin grey soils. These have migrated across the promontory to become cliff-top dunes on the lee coast (Jennings 1967). Groundwater rich in carbonates dissolved from the calcareous dunes has percolated down through joints and down the cliff face, to produce travertine dripstone formations, best seen in the small cave just east of the lighthouse.

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)   

⊡⊡ Fig. 21.6.3.1 Aire River estuary. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.3.2 Landforms of the Cape Otway area. (Courtesy Geostudies.)

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21.6.3

  Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

⊡⊡ Fig. 21.6.3.3 Coastal features between Cape Otway and Seal Point. (Courtesy Geostudies.)

Between Otway Cove and Seal Point (>Fig. 21.6.3.3) the cliffs are cut into Pleistocene dune calcarenite, exposing layers that dip steeply to the SSE, indicating that the dunes were built as sand spilled in this direction across the Cape Otway promontory. At Otway Cove the shore platform is interrupted by a beach of firm, fine sand, and when it resumes to the east the platform is cut into Pleistocene dune calcarenite which here extends down below present low tide level, so that the Cliffs cut in Cretaceous sandstone and mudstone disappear behind steep scrubby slopes on dune calcarenite behind a broad shore platform cut in Cretaceous rocks (Bird 1984). The volution of these coastal features is shown in (>Fig. 21.6.3.4). East of Cape Otway shore platforms are welldeveloped, and show various features. Beyond Point Franklin the steep coastline swings NE, and is more sheltered from the dominant SW ocean swell and storm waves. Broad, undulating shore platforms of landward-dipping Cretaceous sandstone and mudstone stand in front of a steep forested coastal escarpment that extends in places down to high tide level, interrupted by occasional deeplyincised valleys. Some of the smaller streams have valley mouths truncated by cliff recession and arrive at the shore by way of coastal cascades. Elliott River descends steeply from a rocky gorge deeply incised into the Otway Ranges, and flows out over waterfalls into a small cove where calcareous sand has been washed into a gully that interrupts the shore platform. At Storm Point the steep coastal escarpment recedes behind a basal terrace 5–8 m above sea level, which extends

NE for about 2 km past Swell Point to Marengo, where it is truncated by low cliffs and bordered by shore platforms cut in the Cretaceous sedimentary rocks (>Fig. 21.6.3.5). Mounts Bay has a sandy surf beach, backed by grassy and scrubby dunes, facing east-south-east, and receiving ocean swell that has been refracted round Cape Otway and past Marengo. The dunes are backed by the alluvial valley floor of the Barham River, the mouth of which has been deflected northwards behind a wide sand spit with low hummocky dunes, and reaches the sea beside low bluffs. South of Apollo Bay town the riverside bluffs pass laterally into low sea cliffs which are mantled by Holocene dunes, while the shore platform at the river mouth continues, partly obscured by a sandy beach, out to Point Bunbury. The harbour at Apollo Bay (>Fig. 21.6.3.6) was formed when breakwaters were built in the 1950s. These have modified patterns of longshore drifting. Sand began to accumulate on Apollo Bay beach, and as the beach widened the backshore dunes grew higher (Bird and Jones 1988). Since 1955 it has been necessary to dredge sand from the harbour, excavated by suction and piped on to Apollo Bay beach. The beach curves round to Wild Dog Creek, and is backed by grassy and scrubby Holocene dunes and low bluffs marking a former cliffed coastline. Round the bay the dunes diminish, and erosion is severe during episodes of strong easterly wind, when storm waves break heavily on this shore. The scarp that runs behind Apollo Bay is fronted by a broad terrace, and both are breached by the Wild Dog Creek valley, which winds through a deep gorge and opens

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)   

⊡⊡ Fig. 21.6.3.4 Evolution of the coast near Cape Otway: A, in Pleistocene times, with the sea at or a little above its present level; B, during the Late Pleistocene low sea level phase, when calcareous dunes spilled across Cape Otway, burying the cliff and shore platform: these dunes then became hardened into Pleistocene dune calcarenite; C, in Holocene times, after the sea had risen to its present level, cliffs and shore platforms have been cut into the Pleistocene dune calcarenite. Locally the dune calcarenite has been removed and the cliffs and shore platforms are cut in the underlying Cretaceous sandstones. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.3.5 Cliff and shore platform weathering on Cretaceous sandstone at Marengo Point, near Apollo Bay. (Courtesy Geostudies.)

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21.6.3

  Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

⊡⊡ Fig. 21.6.3.6 An aerial view of Apollo Bay harbour, showing swell breaking on Point Bunbury refracted round the breakwater to move in to the prograded beach. The grassy dunes formed after the harbour was built. (Courtesy Neville Rosengren.)

on to the sandy beach. This stream is usually diverted well to the east by a sand spit, towards low cliffs where the shore platform resumes. Segments of shore platform continue eastward to the sandy cove at the mouth of Skenes Creek. The seaward dip in the shore platform is marked by in-facing cuestas on sandstone outcrops and corridors of planation on the mudstones (>Fig. 21.6.3.7), while the narrow beach to the rear runs out in miniature cusps and tombolos in the lee of upstanding mesas of sandstone. The scarp that runs eastward from Apollo Bay comes to the coast at Cape Patton. Sugarloaf Hill is a knoll of Cretaceous sandstone that was an island surrounded by a shore platform when the sea stood a few metres higher in the Late Pleistocene. Where massive sandstones dip landward the steep coast is a bold escarpment, and the shore platform is narrow. Intervening sectors, where the dip is seaward, are generally lower and gentler, with wider shore platforms. Where the dip is landward these have bold outer rims. Where the strike runs at an angle to the shore, as at Carisbrook Creek, sandstone cuestas face alongshore between clefts cut out in mudstones. Sandy beaches occur in successive bays at the Kennett and Wye Rivers, where the southerly ocean swell, already refracted as it moves in towards the coast, curves sharply round headlands and breaks laterally as it moves in towards the beach. The lateral breaks provide good opportunities for surfing. Cannon-ball concretions occur at several places along  the coast. They are slightly harder spheroidal or

­ illow-shaped formations produced within the Cretaceous p rocks as the result of localised cementation of the sandstones by calcite and iron compounds, and they now stand out as the result of differential weathering. They are well displayed at Artillery Rocks (>Fig. 21.6.3.8). Cumberland River and St. George River wind out of gorges, past vertical river cliffs, into sandy coves that interrupt the shore platform. The beaches are of yellow-brown calcareous, washed in from the sea floor: there is very little grey sand from the Cretaceous outcrops along the coast, or from the rivers. The Great Ocean Road between Apollo Bay and Lorne (>Fig. 21.6.3.9) has been disrupted from time to time by landslides, especially where the Cretaceous rocks dip seaward. Excavation of the steep coastal slope to construct a corniche road in the 1920s and 1930s increased instability, and led to the dumping of heaps of angular boulders at the foot of the coastal slope. There has been further disruption as the road has been widened. At Point Grey the coastline turns NW along the strike of the Cretaceous rocks, and there are shore ramps sloping with the north-easterly dip into the Lorne syncline along the sides of Loutit Bay. Waves rounding Point Grey arrive obliquely, sweeping sand along this shore towards the beach at the seaside resort of Lorne. Erskine River, a bouldery watercourse, has an outlet diverted eastward by a sand spit produced by longshore drifting where waves run in obliquely to the shore. East of Lorne the steep coast resumes, with basal cliffing and shore platforms cut in Cretaceous rocks interrupted at the

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)   

⊡⊡ Fig. 21.6.3.7 Shore platform east of Apollo Bay, showing truncation of seawarddipping Cretaceous sandstone with small in-facing scarps. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.3.8 Artillery Rocks, formed where calcareous cannon-ball concretions have weathered out from seaward-dipping strata. (Courtesy Geostudies.)

mouths of steep valleys. Cinema Point is a bold cliff that stands beside the deeply incised winding Grassy Creek valley, which includes two steep-sided residual hills resulting from circum-denudation by the river. Heaps of boulders dumped from road works at the back of the shore have been agitated by storm wave action and worn into adjusted, interlocking shapes by storm waves at high tide. These were termed fitting boulders by Hills (1970).

Successive landslides have occurred at Clarkes Slip, on the eastern side of Point Castries, where weathered mudstones dip seaward. Here the coastal outcrops of Cretaceous sandstones and mudstones come to an end as these rocks pass beneath Lower Tertiary formations in the grassy bluffs at Eastern View. The Tertiary formations dip SE off the Otway Ranges, initially at up to 40°, but soon declining to less than 10° as progressively younger strata outcrop NE, between Eastern

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21.6.3

  Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

⊡⊡ Fig. 21.6.3.9 The steep coast of the Otway Ranges south-west of Lorne, with the Great Ocean Road as a corniche. Surf is breaking across the shore platform, submerged at high tide. (Courtesy Neville Rosengren.)

View and Torquay. The strata are gently folded along the coast, with several small anticlinal and synclinal structures superimposed on the SE dip. A wide sandy beach extends from Eastern View to Fairhaven, backed by grassy dunes and heathy bluffs cut into Palaeocene and Eocene formations. There are ferruginous sandstones, carbonaceous clay, sandy clay and seams of brown coal. At Fairhaven the bluffs pass inland beside the Airey’s Inlet valley. A high foredune continues eastward along a barrier spit that deflects the outlet from Pancalick Creek at Airey’s Inlet. The river opens to the sea beside crumbling cliffs of yellow rubbly Point Addis Limestone, of Oligocene age. This gives place to grey basalt in the shore platform beneath the cliffs of Split Point, which expose a dissected Eocene volcano. This is the first coastal outcrop of the early Tertiary Older Volcanics, which are prominent to the east of Port Phillip Bay. Table Rock, to the south, is a planed-off plug of basalt, and Eagle Rock is a large stack rising from the shore platform to the east, in which the volcanic rocks are capped by Port Addis Limestone showing vertical fluting. The Port Addis Limestone then rises in the cliffs, and to the east these give place to more rounded, grassy bluffs on the underlying (Lower Oligocene) Angahook sands and clays, while basal tuffs and agglomerates continue to form the shore platform NE to Urquharts Bluff. To the east Pleistocene dune calcarenite emerges from the backshore dunes to form low cliffs and rocky shore outcrops in a narrow ridge that runs out to Point

Roadknight. In the bay to the north the coast steepens, and there is a recurrently active coastal landslide in the seaward-dipping Angahook silty clay, which outcrops at Melba Parade in a receding upper cliff behind a pedimentlike slope, strewn with buckshot (ironstone) gravel, partly overgrown by vegetation and incised by gulleys, with a lower cliff at its seaward edge dropping sharply to the beach (>Fig. 21.6.3.10). The cliffs decline towards the Anglesea valley, where the river flows out through a shallow estuarine lagoon to reach the coast by way of a gap in the dune fringe. The sandy beach at Anglesea includes sediment similar in colour, texture and mineralogy to the outcrops in the cliffs to the west, and is an example of a beach partly supplied with sand derived from nearby eroding cliffs. East of the mouth of the Anglesea River there are low cliffs, and large blocks of ferruginous sandstone have fallen to the shore. This capping sandstone ascends eastwards and disappears as the cliff crest rises rapidly to about 50 m in Demons Bluff (>Fig. 21.6.3.11). The cliffs are almost vertical to locally overhanging, cut in coherent, homogeneous material that is often damp with seepage from the cliff face. For the next 2 km the vertical cliff shows paler sediment over these darker deposits, the junction between the two ascending until it stands about half-way up the cliff. The colour contrast suggests two separate formations, but in fact the Anglesea beds are a single Eocene formation, the upper part of which has been weathered to light brown by oxidation of the dark grey carbonaceous material which persists below. The

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)   

⊡⊡ Fig. 21.6.3.10 The landslide at Melba Parade, Anglesea. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.3.11 Vertical cliffs in the Anglesea Beds at Demons Bluff. (Courtesy Geostudies.)

junction is unrelated to bedding planes, and simply marks a contrast between the pale weathered and the dark unweathered sediment. The cliff face is dissected by chasms and chimneys etched out along joint planes, and there are subsiding blocks beneath arcuate breakaways at the cliff crest. Talus cones of dark carbonaceous material have accumulated beneath scars where the lower part of the cliff has collapsed at intervals during the past few years, and there are cracks

and crevices in the upper cliff where vertical sheets of paler sandy silt are separating from the cliff face. These disintegrate quickly into fine sediment after they fall to the shore. There is no shore platform, the sedimentary rocks having proved too soft and structureless to sustain one. Instead a wide and very gently sloping sandy beach is exposed at low tide, while at high tide waves wash against the cliff base, and storm waves break heavily, splashing high up the cliff face.

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21.6.3

  Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)

East of Point Grinder the cliff crest is trenched by a gully that opens to a hanging valley. The ground rises and the SE dip has a seaward component. As a result, the cliff runs inland behind a large landslide where weak formations lubricated by groundwater seepage and overloaded by saturation in rainy periods have subsided down the dip. The landslides extend to the western side of Point Addis, a cliffed promontory of Oligocene limestone. To the east is another landslide on the Angahook beds, and a sandy beach curves round to the Jarosite Headland. A steep coast then runs behind Bells Beach, a steep sandy beach behind a bouldery shore in a bay that faces SE, a famous surfing area. Beyond Bells Beach the coast consists of steep bluffs in Point Addis Limestone, with outcrops of the harder layers, some of which form structural benches, as at Rocky Point, behind a shore platform exposed at low tide. Fallen blocks of limestone litter the shore. To the east the cliff profile becomes gentler on calcareous clay, the result of a lateral transition in lithology from Point Addis Limestone to a calcareous clay, the Jan Juc Marl, a change in rock type indicating a change from a marine to an estuarine environment here in Oligocene times. Bird Rock is a low, flat-topped stack with a hard yellow-brown limestone capping over grey shelly marl. Several arcuate scarplets of gently-dipping limestone curve across the shore, tracing out a minor anticline. The cliffs continue with an eastward dip behind sandy Jan Juc Beach and on to Torquay, where Point Danger has a cliff of limestone capped by a scalloped slope cut in brown

c­ racking clay and a wide shore platform. To the north is Yellow Bluff, where gently undulating limestone is overlain by calcareous clays, the highest of the Tertiary formations seen along this coast, and probably of Middle Miocene age. To the east of Torquay a sandy beach is backed by high grassy Holocene dunes underlain by Pleistocene dune calcarenite, which emerges to Point Impossible, beside the outlet from Thompson Creek at Breamlea. The beach then continues, with salients developed on lava flows where black Plio-Pleistocene basalt (Newer Volcanics) from the Mount Duneed volcano reaches the shore at Black Rocks. The coastline then swings ESE along Thirteenth Beach, backed by high dunes over Pleistocene dune calcarenite, which outcrops in shore platform segments that broaden as cliffs of dune calcarenite rise towards Point Flinders (Mount Colite), where the underlying basalt emerges to form a slightly higher shore platform. Lake Connewarre is a shallow lagoon to the north, where the River Barwon flows in through Reedy Lake, now occupied by extensive rush and reed swamp. The Barwon enters Lake Connewarre by way of a small marshy delta, built where reeds have trapped river silt. The estuary of the Lower Barwon winds for nearly 10 km downstream from Lake Connewarre through salt marshes to reach the sea at Barwon Heads. In its lower reaches it is fringed by mangroves. East of the mouth of Barwon River a sandy barrier spit is covered by dunes up to 20 m high with woodland, the beach on its seaward side continuing in a long gentle curve ⊡⊡ Fig. 21.6.3.12 Abrasion ramp cut in Pleistocene dune calcarenite on the ocean shore at Point Lonsdale. (Courtesy Geostudies.)

21.6.3

Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale)   

past Ocean Grove to Point Lonsdale. The beach is fronted by segments of shore platform cut in Pleistocene dune calcarenite, which become more continuous towards Point Lonsdale, where there are low cliffs cut into the calcarenite, fronted by an abrasion ramp and a subhorizontal shore platform exposed at low tide (>Fig. 21.6.3.12).

References Baker G (1950) Geology and physiography of the Moonlight Head district. Proc R Soc Vic 60:17–43

Bird ECF (1984) Dune calcarenite and shore platforms at Cape Otway, Victoria. Victorian Nat 101:74–79 Bird ECF, Jones DJB (1988) The origin of foredunes on the coast of Victoria, Australia. J Coastal Res 4:181–192 Hills ES (1970) Fitting, fretting and imprisoned boulders. Nature 226:345–347 Hills ES (1971) A study of cliffy coastal profiles based on examples in Victoria, Australia. Z Geomorphol 15:137–180 Jennings JN (1967) Cliff-top dunes. Aust Geogr Stud 5:40–49

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21.6.4 Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

1.  Introduction Port Phillip Bay is a spacious embayment with an area of 1,950 sq km, a coastline 256 km long, and a narrow entrance from Bass Strait, Port Phillip Heads, 3.2 km wide at high tide, through which strong tidal currents (up to 4 m/s) maintain a deeply scoured channel known as The Rip. Because it is almost landlocked, Port Phillip Bay is rather like a marine lake. The northern part is a saucer-shaped basin with a maximum depth of 24 m, but the southern part is shallow, with extensive shoals exposed at low tide between deeper tidal channels that converge towards Port Phillip Heads (Keble 1946). Tide ranges are small (less than a metre at springs): high tides arrive in Bass Strait and are transmitted in through Port Phillip Heads, and then up to Williamstown, arriving more than three hours after passing Point Lonsdale Jetty. Ocean swell entering Port Phillip Heads is quickly diffracted and weakened, and away from the entrance waves are generated entirely by winds that blow across the bay. Port Phillip Bay first became a marine embayment when Tertiary formations were uplifted to form the Bellarine Peninsula to the south, along the east coast of the country. It occupies part of a fault-bounded structural depression or sunkland (Keble 1946) between Rowsley Fault, to the west and Selwyn Fault, to the east. Recurrent subsidence has produced a broad lowland, the NW part of which is covered by the plain of Newer (Plio-Pleistocene) Volcanics which extends between Geelong and Melbourne; the SE part, submerged by the sea, became Port Phillip Bay. The present, almost enclosed, embayment did not develop until the Late Pleistocene, when the dune calcarenites of Point Lonsdale and the Nepean Peninsula had accumulated on either side of the entrance. A bay much like the present one existed 80,000–120,000 years ago when the sea stood a little above its present level, but the modern outlines of Port Phillip Bay were established only about 6,000 years ago, in the final stages of the Late

Quaternary marine transgression, when the Holocene sea flooded back into the Port Phillip basin. During low sea level phases of the Pleistocene, the Yarra and other streams draining into Port Phillip Bay met on the emerged sea floor and flowed out through a deep channel and down to the lowered coastline, probably somewhere between Cape Otway and King Island. The entrance to Port Phillip Bay is constricted by the uplifted Bellarine Peninsula, bounded by the Curlewis Monocline and the Bellarine Fault on the western side, and by the Nepean Peninsula, an area of accumulated Pleistocene dunes partly consolidated into dune calcarenite, overlain by Holocene dune formations on the eastern side. In the NE, the coastline is set back along the Beaumaris Monocline, while in the SW Corio Bay and its fringes occupy a subsidiary depression. On the east coast, there are swampy lowlands behind sandy barrier beaches: Carrum Swamp in the low-lying area between the Beaumaris Monocline and Selwyn Fault, and Tootgarook Swamp in a depression north of the Nepean Peninsula and west of Selwyn Fault as it runs down towards Cape Schanck. The present coastline of Port Phillip Bay is varied. Much of the east coast is steep and cliffed, with rocky shores and beaches of quartzose sand and gravel mainly in coves and embayments, but there are longer beaches behind nearshore sandy areas on the northern shore, between the mouth of the Yarra and Beaumaris, on the  eastern shore between Mordialloc and Frankston and on the southern shore between Safety Beach and Sorrento, where the beach sands change from predominantly quartzose to mainly calcareous as they pass cliffed  outcrops of dune calcarenite west of Rye. Intertidal  and nearshore sand bars are well developed in  front of beaches along parts of the east and south coast. The west coast between Geelong and Melbourne borders a low-lying (Newer Volcanics) plain with a few segments of cliff cut into Quaternary sediment and areas of salt marsh, mangroves and alluvium, and shelly beaches

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

along much of the shore, sometimes with irregular intertidal bars of fine calcareous sand. The coast of the Bellarine Peninsula is generally higher and steeper, lined by narrow beaches of sand and gravel. Spits of sand and gravel partly enclose shallow Swan Bay, bordered by wide salt marshes, and Mud Islands lie offshore, surmounting the wide sandy threshold in the southern part of the Bay. Wave action, determined largely by the wind regime over Port Phillip Bay, is stronger on the east coast, exposed to the prevailing westerlies. Nearshore currents are produced by a combination of wind, wave and tide, with the tidal effects increasing towards Port Phillip Heads, where refracted ocean swell moves into Lonsdale Bay and the coast to Observatory Point. On the east coast, the north and NW winds in winter produce waves generating southward drifting of sand along the beaches, while in the summer months waves produced by the dominant south and SW winds reverse this. There are also short-term

a­ lternations related to meteorological conditions. Patterns of sediment transport on the shores of Port Phillip Bay are shown in (>Fig. 21.6.4.1). There are features suggestive of a phase of higher sea level around Port Phillip Bay within Holocene times, probably between 6,000 and 4,000 years ago. Hills (1940) concluded that although some of these features could be the result of local tectonic uplift, the evidence was so widespread that a eustatic fall of sea level had probably contributed to it. Many headlands and shore platforms show induration as the result of precipitation of cementing material in rock cavities. Some of this is due to seepage of groundwater carrying carbonates or iron oxides which are deposited on or near the surface as the rock dries out. Within the past 6,000 years, there has been accretion on parts of the shores of Port Phillip Bay. Beach ­sediment include quartzose sand derived from the weathering of

⊡⊡ Fig. 21.6.4.1 Patterns of beach drifting on the shores of Port Phillip Bay, resulting from wave patterns determined by the summer and winter wind regimes. (Courtesy Geostudies.)

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

cliff, foreshore and sea floor outcrops of sandstone and granodiorite, mainly on the east coast; black sand originating from weathered basalts on the west coast; calcareous sand associated with the dune calcarenites to the south; shelly beaches derived from organisms that live on shore platforms and nearshore reefs, and gravel beaches derived from the coarser and harder components of shore outcrops, notably ferruginous sandstones. Marshlands, confined to sheltered inlets, such as Swan Bay and Limeburners Bay on the west coast, are areas of muddy sedimentation, with salt marsh and some fringing mangroves. In recent decades, the most obvious changes have been the building of sea walls and boulder ramparts to halt cliff recession, especially where it threatened to undermine developed land, and the emplacement of artificial beaches on several sectors. The prediction of global warming leading to a worldwide sea level rise has important consequences for Port Phillip Bay, where coastline erosion would accelerate and low-lying areas would be submerged (Bird 2006).

2.  The Coastline Rugged cliffs of dune calcarenite beneath Point Lonsdale lighthouse overlook flat shore platforms. Surf breaks heavily over outlying calcarenite reefs. When strong SW winds accompany high tides, sand lobes migrate along the beach from Ocean Grove and sand is swept round Point Lonsdale and into Lonsdale Bay. Beach erosion has been a problem at Point Lonsdale for several decades. It has been suggested that recurrent deepening of the entrance channel beneath The Rip has resulted in stronger waves and currents in Lonsdale Bay, but similar erosion has occurred on beaches elsewhere in Port Phillip Bay and along the Victorian coast, and indeed has been prevalent on a global scale over the past century (Bird 1985) so that there is no need to invoke a local explanation for beach erosion. The response to beach erosion at Point Lonsdale has been to construct sea walls, initially timber, then masonry and then boulder ramparts. There has been a ‘domino effect’: after each phase of sea wall construction, beach erosion intensified on the next sector to the north so that successive extensions of the sea wall are set back landward. Beach erosion and dune cliffing continue beyond the northern limit of the last boulder rampart built in 1977 (Bird 1980). The dune calcarenite at Point Lonsdale is backed by diverging ridges that run through to the Port Phillip Bay coast at Queenscliff. These ridges originated as successive

21.6.4

spits built by the longshore drifting of sand from the west during Late Pleistocene times, constricting the entrance to Port Phillip Bay. There are intervening shallow brackish shelly lagoons and salt marshes. Shortland Bluff is a cliff cut into dune calcarenite at the termination of one of the incurving ridges behind Point Lonsdale. Construction of a breakwater to maintain the entrance to Swan Bay Harbour at Queenscliff in the 1930s, with subsequent extensions, has led to the trapping of northwarddrifting sand to form a wide beach and triangular vegetated foreland in front of the masonry wall that had been built to stabilise the Queenscliff seafront. Sand has been pumped through a pipe from the beach on the south side of the breakwater at the ferry terminal across to Swan Island. This led to the growth of a wide recurved spit, known locally as Sand Island, north of the harbour, backed by intertidal mudflats. Swan Island consists of dune ridges running northward on a breached spit formation, and intervening salt marshes. Successive sandy lobes have moved northward along the beach on the east coast of and around the northern end to form a spit that has grown westward. Swan Bay is shallow and fringed by salt marshes, especially on the western shore, which is sheltered from waves generated by the prevailing westerly winds. The marshes are backed by low bluffs marking the early Holocene coastline formed about 6,000 years ago. Edwards Point is a multiple spit of sand, shells and ferruginous gravel, which has grown south from the cliffs at St. Leonards. Indented Head, a low cliff of ferruginous sandstone backs a gravelly foreshore. It marks a division between southward drifting of beach sediment to Edwards Point and westward drifting towards Point Richards, with the divergence produced by waves arriving from the NE. The low-lying coast continues past Point George and Grassy Point, lobate promontories fronted by sandstone reefs dissected into a gravelly shore, with nearshore sand bars and intertidal sand flats carpeted with seagrass, until the higher ground that runs behind Portarlington intersects the coastline. The cliffs become bluffs at Portarlington and run inland behind the large triangular Point Richards foreland, which consists of subdued sandy beach ridges with shell beds separated by swales and backed by an elongated swamp. The bluffs return to the coast to the west, north of the village of Bellarine, and continue along the Curlewis Monocline, parallel to the coast, as a series of cliffs up to 15 m high, with minor headlands and several active landslides. Near Leopold the coastal bluff declines across an N-S fault and swings inland. Beaches and beach ridges of shelly

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21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

sand curve out to a cuspate spit at Point Henry. The bay to the west is backed by the salt marshes of Stingaree Bay, now an area of evaporation ponds where salt is produced. At Geelong the shore of Corio Bay is largely artificial, bordered by sea walls, jetties and a marine swimming pool and then wharves. To the north, the beach diverges from the coast as a shelly sand spit, one of the two spits that constrict the mouth of Limeburners Bay, an estuary fringed by mangroves and salt marsh in an incised valley at the mouth of Hovells Creek. On the northern shore of Corio Bay, beaches of white shelly sand near Avalon alternate with promontories and reefs of basalt, as at Point Lillias, which is fringed by boulders of black lava. This marks the beginning of the lowlying plain of Plio-Pleistocene Newer Volcanics, which borders the west coast of Port Phillip Bay between here and the mouth of the Yarra River at Williamstown. The Sand Hummocks are low barrier spits of shelly sand that partly enclose a shallow tidal lagoon, backed by scrubby salt marsh. Between this and the mouth of the Werribee River the coast is largely artificial, consisting of embankments and sluices bordering the aeration ponds and pastures of the Werribee sewerage farm, which treats waste from Melbourne and uses it to irrigate pastureland. The Little River, The Werribee River and Kororoit Creek each cross the Newer Volcanics plain in steep-sided valleys that were incised during the Late Pleistocene low sea level phase and partly infilled by Holocene sedimentation during and since the ensuing marine transgression. In the Pleistocene, the Werribee River built an extensive deltaic plain of silty alluvium into which the modern valley has been cut, widening downstream from the town of Werribee and bordered by cliffs exposing brown silty Pleistocene alluvium. North of the Werribee delta the sandy beach broadens towards Point Cook, a lobate foreland consisting of a succession of roughly parallel sand ridges, backed by lagoons and salt marshes. The coastline curves to face south at Altona, where the sinuous drift-dominated beach was artificially renourished in 1982. The Newer Volcanic basalt emerges at Williamstown to form a series of wide ledges along a boulder-strewn shore platform, with volcanic structures, such as lava blisters, well exposed. The coastline becomes artificial, lined by sea walls and piers, towards the mouth of Yarra River, and the Melbourne docks. The Yarra, joined by the Maribyrnong River has a bay-head delta south of the city of Melbourne. The coast curves south-east past Station Pier at Port Melbourne and along the sandy beach to St. Kilda. Elwood Beach occupies the first of a series of bays, separated by rocky

promontories, on the coast that borders the south-­eastern suburbs of Melbourne down to Beaumaris. The coast has been much modified by sea wall construction, especially since the 1930s. It is formed on gently undulating Tertiary formations, the soft Red Bluff Sand resting upon harder, more ferruginous Black Rock Sandstone. Cliffed promontories such as Green Point, Picnic Point and Red Bluff occur where the harder Black Rock Sandstone rises above sea level, while intervening bays have been cut in softer Red Bluff Sand at Brighton, Hampton and Sandringham, where sandy beaches are backed by low scrubby bluffs. Offshore breakwaters, linked to the land by piers, have been built to shelter the harbours at St. Kilda and Middle Brighton. Converging wave action in the lee of the breakwaters has formed cuspate forelands bearing low hummocky dunes, and there has been a depletion of beaches to the north and south (>Fig. 21.6.4.2). The beaches on Melbourne’s bayside coast show seasonal alternations in longshore drifting. In the summer months, when southerly and south-westerly waves predominate, sand moves northward along the shore, widening the beaches at their northern ends. In winter, the waves arrive mainly from the west or north-west and move the sand back southward. These seasonal alternations of beach drifting have caused problems, as in the Hampton Beach compartment between Green Point and Picnic Point. Construction of a breakwater to shelter Sandringham Harbour created a trap for sand drifting southward each winter, the breakwater excluding the southerly wave action that previously returned it northward in summer. As a result, Hampton lost most of its beach, and Sandringham Harbour was reduced in area and shallowed by the accumulation of sand. Subsequently, Hampton Beach has been artificially renourished, with groynes designed to prevent sand drifting south into the harbour. South of Picnic Point, Sandringham Beach occupies another embayment, backed by scrubby bluffs. Here, too, northward drifting of sand widens the beach near Picnic Point in the summer months (November–April), and there is southward drifting towards Red Bluff during the winter. The beach has been gradually depleted in recent decades, and in the southern part the beach was very narrow in winter, the backing bluffs were undercut by wave erosion, forming slumping cliffs. The beach has been artificially renourished to prevent further cliff erosion. Red Bluff (>Fig. 21.6.4.3), the highest cliff bordering Port Phillip Bay (34 m), is the geological type section for the Red Bluff Sand and Black Rock Sandstone. Half Moon Bay in a cove between Red Bluff and Black Rock Point is

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

21.6.4

⊡⊡ Fig. 21.6.4.2 Middle Brighton harbour, showing cuspate foreland in the lee of the offshore breakwater. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.4.3 The headland at Red Bluff, Sandringham shows soft, gullied Red Bluff Sands overlying harder Black Rock Sandstone which protrudes as a dark basal ledge. (Courtesy Geostudies.)

s­ hallow and backed by a sandy beach in front of steep scrubby bluffs. The battleship Cerberus was grounded offshore here in 1926 to act as a breakwater protecting a small boat harbour, but it is rusting and slowly disintegrating. The rilled cliffs at Black Rock Point (>Fig. 21.6.4.4) have been cut in pale Red Bluff Sand overlying darker Black Rock Sandstone, which protrudes to form a structural shore platform. The Black Rock Sandstone is light brown

at the cliff base, darkening and hardening seaward across the shore platform, which is littered by blocks broken off and thrown up by storm waves. Potholes have been carved out by waves circulating lumps of hardened sandstone in the softer mid-platform rock. South from Black Rock Point, the beach at Black Rock also shows seasonal alternations related to wave incidence (>Fig. 21.6.4.5): the northern end of the beach becomes

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21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

⊡⊡ Fig. 21.6.4.4 At Black Rock Point rilled cliffs in Red Bluff Sand back a shore platform in Black Rock Sandstone overwashed by waves at high tide. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.4.5 Landforms on the coast at Black Rock, showing wave roses (determined by wind regimes: percentage frequencies indicated) for nearshore waters. These result in a dominance of northward drifting of beach material towards Black Rock Point each summer and southward drifting towards Quiet Corner each winter. (Courtesy Geostudies.)

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

21.6.4

⊡⊡ Fig. 21.6.4.6 The northern part of the beach at Black Rock in late summer, after it has been widened by northward drifting. It is backed by thickly-vegetated coastal bluffs. (Courtesy Geostudies.)

wider during summer (November–April) (>Fig. 21.6.4.6) and narrower during winter. Seasonal drifting results in alternating patterns of accretion alongside a mid-beach groyne (>Fig. 21.6.4.7). At the southern end of Black Rock Beach, the two promontories at Quiet Corner are lined by a sea wall and an undercliff wall built in the late 1930s to stabilise vertical receding cliffs as artificial grassy bluffs graded to a slope of about 30°. As a result, the beach fronting the sea wall was diminished by wave reflection. At Quiet Corner and Beaumaris there are intertidal seagrass terraces, often edged by clifflets, dissected by rounded ponds (scoured by currents circulating as waves move over at high tide), some of which have been opened as rounded bays. Between Quiet Corner and Beaumaris the shore platform exposed at low tide is dominated by gently undulating layers of rock that have been indurated by ferruginous precipitation. These are being cut back at the outer edge and dissected by abrasion along fissures as waves move stones to and fro, and potholing as waves circulate stony material in hollows. At Table Rock Point the coast swings ENE along the axis of a monocline. Steep cliffs of Black Rock Sandstone up to 12 m high line the shore of Beaumaris Bay, with only minor beaches. The sandstone is well jointed, and occasional rock falls have produced heaps of blocky talus which have hardened on exposure to protect parts of the cliff base from wave attack.

On the plateau hinterland between Sandringham and Mentone, the suburban landscape is hummocky, the Tertiary sandstones and clays being overlain by ridges of quartzose dune sand trending NW–SE. When European settlers arrived in the 1830s, these subdued dune ridges were stable beneath a cover of heath and scrub woodland. They evidently formed during a Pleistocene arid phase (Whincup 1944). At Mentone the coast resumes its south-eastward trend, and the former cliffs of sandy clay have been stabilised as scrubby slopes behind a concrete sea wall. Here, as elsewhere, the sea wall caused wave reflection which scoured away the beach, and was artificially replaced in 1977 and renourished in 1984. Some of the sand eroded from Mentone Beach has drifted southward to Mordialloc, where a beach 100 m wide has formed alongside the breakwaters built to stabilise the mouth of Mordialloc Creek. On the southern side of the breakwaters, the sandy coastline has been cut back. Between Mordialloc and Frankston the quartzose sandy beach is long and gently curved, its shape determined by the refraction of waves produced by the prevailing westerly winds over Port Phillip Bay as they move in through gradually shallowing nearshore water. The beach has a slightly sinuous outline, with lobes that move to and fro along the shore when oblique waves generate northward or southward drifting. It is fronted by two or three parallel nearshore sand bars, with gaps through which rip

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21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

⊡⊡ Fig. 21.6.4.7 Sand accretion alongside a concrete pipe in February (above) indicates northward drifting in summer and in September (below), southward drifting in winter. (Courtesy Geostudies.)

currents flow. The sand is fine, as the result of attrition during wave-reworking, and the beach has been eroding in recent decades, but there are usually splays of unwashed shells. Seaford beach forms the seaward margin of a sandy barrier system on which foredunes mark stages in progradation. It is an outer barrier backed by Carrum Swamp, a former lagoon, then by an inner (Pleistocene) barrier, behind which is the large Dandenong Swamp. The ­coastline has been relatively stable, apart from some minor

backshore cliffing of the dunes. Beach sand moves northward in summer and southward in winter, as shown by seasonal variations in the pattern of accretion alongside the concrete drain near the Riviera Hotel, Seaford. The sandy beach comes to an end beneath Olivers Hill, where Selwyn Fault brings up the Tertiary and older rock formations at the beginning of a cliffed coastline. There are outcrops of much-weathered Mount Eliza ­granodiorite and ferruginous Baxter Sandstone (Lower Miocene). To the south are slumping clay cliffs, then a series of sandy

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

bays between rocky headlands with shore platforms exposed at low tide, past Mornington to Mount Martha. Along Sunnyside Beach there are intertidal outcrops of several formations brought up to the south of the Manyung Fault, including basalt (Older Volcanics), granodiorite and a small outcrop of pale grey Cretaceous mudstone, with some lignite seams. The scrub-covered backing slope is disrupted by landslides with hummocky slumped topography on Balcombe Clay (Upper Miocene) below an upper breakaway. To the south, granodiorite forms steep coastal bluffs and a rocky, boulder-strewn foreshore, with pocket beaches of locally-derived cobbles and pebbles and some remnants of a pebbly raised beach about a metre above present high tide level. To the south the rugged granodiorite shore comes to an end against dark brown Baxter Sandstone, which forms large boulders on the shore as sandy beaches resume. In the Mornington area, sandy beaches occupy coves separated by cliffed outcrops of Baxter Sandstone, the most prominent being Schnapper Point and Linley Point, both with emerged shore platforms, probably of Holocene age, near the base of the cliffs. A breakwater shelters Mornington Harbour, and to the south the Baxter Sandstone forms a succession of high promontories and cliffs, with boulder beaches of fallen sandstone hardened into dark brown ironstone. The underlying Balcombe Clay then appears at the base of the cliffs, and in Balcombe Bay there have been landslides where Baxter Sandstone has subsided over slippery Balcombe Clay to form irregular tumbled coastal slopes. ⊡⊡ Fig. 21.6.4.8 The cliff and shore platform cut in weathered granodiorite on the southern coast of Mount Martha. (Courtesy Geostudies.)

21.6.4

Balcombe Bay beach ends against cliffs of weathered granodiorite, which have a shore platform consisting of a wave-cut (abrasion) ramp. The superficial weathered zone of the granodiorite rises southward with the contours of the Mount Martha upland, so that beyond Balcombe Point the cliffs become steeper on the underlying harder unweathered rock and the abrasion ramp disappears (Jutson 1940). The rocky coast continues past Martha Point where the steep slopes decline to low cliffs, and the bordering abrasion ramps reappear as the more weathered granodiorite zone descends again to sea level (>Fig. 21.6.4.8). There are cobble and pebble beaches derived from the adjacent cliffs and shore platforms where the rocks are closely jointed and weather into small angular fragments. These are added to the beach and rounded by attrition as the waves agitate them. The southern part of the Mornington Peninsula has a NW projection, known as the Nepean Peninsula, which narrows towards Point Nepean at Port Phillip Heads. Safety Beach marks the beginning of a long sandy beach that extends SW, past Dromana and Rosebud, to Rye and Blairgowrie. Initially, the beach sand is quartzose and includes minerals derived from the Mount Martha granodiorite, but beyond Dromana it becomes more calcareous, passing into the province of the Pleistocene dune calcarenites of the Nepean Peninsula. The coast skirts the steep slopes of Arthurs Seat, scarcely interrupting the beach. Dromana Pier is one of several structures which carry cushions of the tubeworm Galeolaria caespitosa, the upper limit of which stands about 30 cm above mid-tide level

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21.6.4

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

⊡⊡ Fig. 21.6.4.9 The tubeworm Galeolaria caespitosa on pier supports at Dromana is zoned in relation to tide levels, the upper limit being almost horizontal, just above mid-tide level. (Courtesy Geostudies.)

and could be monitored to provide an indicator of sea level change (>Fig. 21.6.4.9). Between Dromana and Blairgowrie the beach is backed by a barrier of low sandy beach ridges bearing dune ­woodland. The intertidal zone widens and the ebbing tide exposes up to ten sand bars with intervening swales, generally parallel to the coastline (>Fig. 21.6.4.10). These show minor short-term alternations in morphology and configuration during stormy weather when large waves break over them, but the nearshore topography has remained generally stable. At Rosebud the sand bars have been modified by the dredging of a boat harbour, and attempts were made to augment the beach by bulldozing sand from the intertidal zone. However, the scrapings were colonised by seagrasses (Zostera sp.), and large quantities of seagrass hay began to accumulate on the beach. The problem diminished after sand was deposited at right angles to the shore, to be dispersed laterally by wave action to blanket and destroy the seagrass beds. At Rye the beach was artificially nourished in 1972, and there has been some land reclamation. White Cliffs to the west and The Sisters near Sorrento, are headlands of the Pleistocene dune calcarenite that forms the foundation of the Nepean Peninsula, and may represent terminations of recurved ridges matching those seen in the Point Lonsdale area, a spit formation buried by extensive surmounting dunes. Between Sorrento and Portsea there are more cliffs and bluffs of dune calcarenite, with headlands where the more resistant calcrete layers outcrop at sea level, as at Point King, and an intermittent, narrow and

⊡⊡ Fig. 21.6.4.10 Multiple sand bars in the intertidal zone at Blairgowrie, in the SE of Port Phillip Bay. (Courtesy Geostudies.)

­ issected shore platform. Point King Beach, to the west, d shows westward longshore drifting by waves from the NE and eastward drifting by waves from the NW, but is also influenced by oblique waves produced by ferries plying between Sorrento and Queenscliff.

Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean)

21.6.4

⊡⊡ Fig. 21.6.4.11 Mud Islands, Port Phillip Bay. (Courtesy Geostudies.)

To the west of Portsea there are intermittent cliffs cut across the dune calcarenite ridges of the Nepean Peninsula, with beaches in intervening coves. The beach curves out to Observatory Point, a large sandy lobate foreland bearing successively-built foredunes in front of bluffs marking an earlier cliffed coastline cut in dune calcarenite. This beach is unusual in that it receives surf produced by ocean swell. To the west are cliffs and slopes of Pleistocene dune calcarenite and unconsolidated Holocene dune sand. The cliffs are unstable where the dune bedding dips seaward, as at Weeroona Point. The beach is narrow and transitory, and erosion has been a problem here, resulting in the building and repair of sea walls at intervals over the past century. The possibility that the sea would break through and form a second marine entrance to Port Phillip Bay has been countered by building a boulder rampart along the narrowest section, which on the Bass Strait side is fortified by a sea wall behind a shore platform cut in dune calcarenite. At Point Nepean the sea wall has stabilised former cliffs, but a broad shore platform is exposed at low tide when it is possible to walk out to the rocky stacks that border The Rip. Some oceanic sand, including shelly material and fragments of calcarenite, has been washed up from the sea floor on to low sectors of the shore platform, and thence to the beach. Out in Port Phillip Bay are Mud Islands (>Fig. 21.6.4.11), a group of sandy barriers bearing vegetated dunes that rise

up to 3.6 m above high spring tide level. Shaped like an emerged atoll, these barriers enclose a shallow lagoon, with a floor of sandy mud, bordered by salt marsh. The sand has been washed up from the surrounding shoals, and is mainly calcareous, with some quartz and fragments of shelly material. Mud Islands were first mapped in 1836, and their portrayal on subsequent maps, charts and air photographs shows that their configuration has varied as the result of erosion and deposition, with phases when neighbouring shoals of sand were washed onshore. At present there are two inlets through which sea water enters and leaves the lagoon as the tides rise and fall, and both have in-growing sandy tidal deltas.

References Bird ECF (1980) Historical changes on sandy shorelines in Victoria. Proc R Soc Vic 91:17–32 Bird ECF (1985) Coastline changes: a global review, Wiley, Chichester Bird ECF (2006) The effects of a higher sea level on the coast of Port Phillip Bay. Vic Nat 123:49–54 Hills ES (1940) The question of recent emergence on the shores of Port Phillip Bay. P R Soc Vic 52:84–105 Jutson JT (1940) The shore platforms of Mount Martha, Victoria. Proc R Soc Vic 52:164–174 Keble RA (1946) The sunklands of Port Phillip Bay and Bass Strait. Mem Nat Mus Vic 14:69–122 Whincup S (1944) Superficial sand deposits between Brighton and Frankston, Victoria. Proc R Soc Vic 56:53–76

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21.6.5 Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

1.  Introduction The Nepean Peninsula consists of dune topography, with Pleistocene dune calcarenites overlain by unconsolidated Holocene dunes along the seaward fringe (Bird 1982). The Pleistocene and Holocene dune formations consist mainly of calcareous sand, with varying proportions (usually less than 20%) of quartz. The sand is biogenic, containing shell fragments and marine organisms that have come from the continental shelf, where similar sediment are forming at the present time. A borehole drilled at Sorrento in 1910 showed that dune calcarenites extend at least 140 m below sea level, with interbedded layers of marine sediment,

including shelly beach sands and marine clays of the kind that are now accumulating on the floor of Port Phillip Bay. The rock sequence indicates several phases of dune sand accumulation interrupted by marine submergence and deposition during the intermittent subsidence of the country to the west of Selwyn Fault, which runs across to Cape Schanck, a promontory of uplifted Older Volcanics (>Fig. 21.6.5.1). Large quantities of biogenic calcareous sand have moved in from the sea floor during Pleistocene times, particularly during low sea level phases when onshore sand collected this sediment from the emerged continental shelf and swept it landward. Some sand may also have

⊡⊡ Fig. 21.6.5.1 Geology of the Nepean Peninsula. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.5, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

been washed up on to the beaches and blown onshore during periods when the sea was rising or falling. The Late Quaternary evolution of the Nepean ocean coast is summarised in (>Fig. 21.6.5.2). As the sea reached its present level about 6,000 years ago it cut away the Holocene dune fringe and began to shape cliffs and shore platforms cut into the underlying Pleistocene dune calcarenite. The Holocene dunes thus persist as cliff-top formations, which in places are still spilling landward. The indented southern coast of the Nepean Peninsula consists of rugged cliffs up to 50 m high, fronted by almost

horizontal shore platforms of Pleistocene dune calcarenite. They are interrupted by sandy surf beaches, backed by Holocene dunes, at Portsea, Rye and Gunnamatta. The cliffs expose sections in the Pleistocene dune calcarenite, some with thin seams of biscuit-like rock inclined at about 30° in the cliff face (>Fig. 21.6.5.3), formed by superficial cementation of sand during phases of stability on the slopes of an intermittently advancing dune. There are interbedded layers of hard calcrete formed as the result of precipitation of carbonates from groundwater raised to surface or subsurface horizons. The palaeosols that overlie

⊡⊡ Fig. 21.6.5.2 Evolution of the Nepean ocean coast: A, About 80,000 years ago, when sea level was falling, multiple dune ridges formed on the emerging sea floor; B, About 10,000 years ago, when sea level was rising, these dunes were eroded by waves, and some sand was blown landward, up and over earlier Pleistocene dune calcarenites; C, The margin of the Pleistocene dune calcarenite has been cut back to form shore platforms and cliffs capped by Holocene dunes. (Courtesy Geostudies.)

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

21 .6 .5

⊡⊡ Fig. 21.6.5.3 Dune sandstone with steeply inclined bedding indicating dune advance over a buried soil (palaeosol) containing root structures ­(rhizoconcretions). (Courtesy Geostudies.)

these calcretes include thin dark organic rendzinas, and often thicker reddish-brown terra rossa. Traced laterally, these soil and calcrete layers vary in thickness, and rise and fall as they mark out the contours of successivelyformed dune landscapes. During periods of topographic stability the dunes carried vegetation that is commemorated by root concretions (rhizoconcretions) beneath the palaeosols and calcrete horizons. The buried Pleistocene soils are sandy, with varying proportions of red or brown clay. Some of this material (notably quartzose sand) may be insoluble residues that accumulated on the surface as the underlying dune calcarenites were dissolved away by rain water, but the bulk of it arrived as wind-blown dust, accumulating to thicken the soil. The palaeosols branch and re-unite in a way that indicates that there were phases of local instability when blowouts developed within the generally stabilised dune landscape and spilled sand downwind. This process continues on top of the cliffs, where unconsolidated Holocene dunes are locally spilling inland to bury the vegetation and soils on the Pleistocene dune topography. Blowouts in these younger cliff-top dunes expose dark coloured buried soils, and organic matter extracted from one of these gave a radiocarbon age of about 5,350 years bp, indicating that the cliff-top dunes are of Holocene age. The landscape of the Nepean Peninsula is dominated by hummocky dune terrain, largely stable beneath a cover of scrub, woodland, or pasture. The dune crests generally reach 20–30 m above sea level, but some rise higher, with

summits locally exceeding 60 m in the Sorrento area. There are no surface streams or lakes for the dune sands are highly permeable and rainfall quickly percolates into the ground, to emerge as springs. Carbonates dissolved in the groundwater are precipitated where the water seeps out of cliffs or drops from the roofs of caves to form layers or sheets of travertine. The outlines of the cliffs along the Nepean Ocean Coast are related to variations in structure and resistance of the Pleistocene dune formations. Rocky promontories run out at intervals along the coast, and some have become detached as stacks which rise from the shore platform (>Fig. 21.6.5.4). Steeply inclined and vertical fissures in the dune calcarenite have been excavated by waves and wind to form clefts and caves, and eventually arches through headlands. At Portsea London Bridge is an islet which has been penetrated by an arch, cut out by waves in the layered dune sandstones beneath a harder calcrete layer, which forms the roof. A former arch near the southern end of St. Pauls Road has collapsed, leaving a stack separated from the cliffs by a heap of tumbled boulders. Headlands are found where the more resistant calcrete layers outcrop at sea level, and these layers also form structural ledges on cliff faces and around stacks, as at Diamond Bay, where storm waves have washed away overlying softer material. These ledges have formed with the sea at its present level, and cannot be taken as indications of emergence. The cliffs are steep, often vertical, and at several places

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21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

⊡⊡ Fig. 21.6.5.4 Rock stacks cut in dune calcarenite on the shore platform near Sorrento. The irregular outlines have been produced by solution of the rock in rainfall, surf and sea spray. (Courtesy Geostudies.)

downwashed dune sand has been cemented to the cliff face as sand stalactites. Along the ocean coast cliffs are receding as they are attacked and undermined by storm waves. Cliff recession has been more rapid on sectors where the shore platform is narrow, or absent, so that storm waves reach the back of the shore without dissipating their energy in breaking across the platform. Subjected to wave attack, dune calcarenites disintegrate to sand, with pebbles and boulders derived from the more resistant calcrete layers. At the cliff base there is often a sloping ramp formed by abrasion, as waves move sand and gravel to and fro. This declines to about mid-tide level, and is fronted by a shore platform which is horizontal, or has a very gentle seaward slope. The shore platform is up to 100 m wide, and is best developed where the dune sandstones are homogenous and of moderate resistance: where they are harder they persist as structural reefs and ledges, and where they are softer they have been removed by marine erosion. The evolution of the shore platform has proceeded in two stages. As the cliffs were cut back abrasion ramps were formed at their base by wave scour: these were then lowered and flattened to about mid-tide level by subaerial weathering processes, notably solution in rain water and sea spray, until they formed horizontal shore platforms exposed during low tides. The bedding of the dune sandstones can be seen truncated in the shore platform, while the more resistant calcretes form projecting outcrops, as at Koreen Point. The shore platforms thus result from a

combination of high wave energy to cut back the cliffs and develop an abrasion ramp, and a moderate tide range, allowing weathering processes to operate during low tide. These weathering processes can be seen at work on ridges and stacks of weathered dune calcarenite in front of the cliffs and rising as residual formations on the shore platform. There are pinnacles of intricately dissected dune rock (>Fig. 21.6.5.5), which owe their form to corrosion by rain water and sea spray, and to the scraping, drilling and plucking effects of the various marine organisms, notably periwinkles, barnacles and mussels, which live on the pitted and honeycombed surfaces. The shore platform is thus produced by the removal of rock material down to a specific level. Hills (1971) postulated that this level was determined by an interaction between seeping groundwater (percolating rainfall), which has dissolved calcium carbonate, and sea water already saturated with calcium carbonate, resulting in precipitation of cementing material in cavitites within the dune rock at mid-tide level. The shore platform is thus a relatively durable ledge, being undermined by wave attack along the scalloped outer edge, which falls away abruptly to depths of at least 5 m, and is typically draped with kelp and other brown seaweeds (>Fig. 21.6.5.6). Potholes of various kinds are seen on these shore ­platforms. Some have been drilled out by waves swirling stones trapped in a cleft or depression which may grow to  depths of 2 or 3 m; others result from solution processes which develop saucer-shaped basins with pitted and

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

21 .6 .5

⊡⊡ Fig. 21.6.5.5 Pitting and spikes (karstic weathering) on dune sandstone on the shore near Jubilee Point, Nepean Ocean Coast. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.5.6 Collapsed arcs on the outer edge of the shore platform at Sorrento. There is adhering kelp below the low tide line. (Courtesy Geostudies.)

­ verhanging rims; still others are due to the excavation of o soil material from case-hardened cylindrical soil pipes that penetrate downward from palaeosols into the underlying calcrete. Basal erosion on headlands and around stacks has formed notches close to mid tide level with overhanging visors about 2 m higher. The rock surface within notches and visors is usually intricately pitted, showing that solution by sea spray has been a contributory process.

Jubilee Point (>Fig. 21.6.5.7) is a large promontory capped by a calcrete layer, with overlying Holocene dunes that have been trimmed back from the cliff crest. The relatively soft underlying bedded dune sands have been etched out by wind scour to form caves high in the cliff face, beneath layers of calcrete. Beneath the headland a large cave has been excavated by storm waves along a crack in the calcarenite rocks to penetrate the relatively hard rocks in the cliff face and scour away the softer backing sandstone.

1353

1354

21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

⊡⊡ Fig. 21.6.5.7 Cliff and shore platform cut in Pleistocene dune calcarenite at Jubilee Point, showing soft dune sandstones with horizons of calcrete and associated palaeosols (arrowed), with a capping of unconsolidated dune sand. (Courtesy Geostudies.)

Beaches on the ocean coast of the Nepean Peninsula occupy coves and embayments, especially where there are gaps in the shore platform, as at Portsea and Rye. The beach sediment are partly of local derivation, consisting of calcareous sand and calcrete pebbles eroded from dune calcarenite outcrops in nearby cliffs and shore platforms, but sand has also been washed in from the sea floor. In some places, as at London Bridge and St. Pauls Cove, the beach consists largely of coarse, well-rounded quartz sand. As quartz sand forms only a small proportion of the adjacent rock outcrops, these coarser beaches appear to have been derived from a mixture of quartz and carbonate sand washed up from the sea floor through gaps and clefts in the shore platform. Wave swash, spray and rainwater have gradually dissolved away much of the carbonate, leaving insoluble quartz sand. The beaches are shaped by ocean swell, refracted as it moves across the shore platforms at high tide, the surf being stronger where there are gaps in the shore plat­­ form. They are subject to cut and fill as storm waves alternate with constructive swell, and sand is also moved along them by waves arriving obliquely, and by longshore winds. To the east of Rye the cliffs and shore platforms are intermittent, and there are long stretches of sandy beach backed by dunes, partly stabilised by grasses and scrub, and partly active and mobile, as at Gunnamatta. Dune calcarenite cliffs and shore platforms with scattered stacks reappear to the south-east, and at Fingal Beach a 40-m high cliff shows a sequence of dune sandstones, calcretes, palaeosols and gravelly breccias.

The cliffs pass into steep vegetated calcarenite bluffs along Fingal Beach, and towards its southern end the underlying Older Volcanics rise to the surface to outcrop in the boulder-strewn shore platform. A cliff section shows a slumping brown clay derived from subaerially weathered basalt, overlain by Pleisto­­ cene dune calcarenite. The well-jointed black basalt disintegrates into angular boulders, which are mixed with the fallen dune calcarenite blocks that litter the beach. Agitated by waves, the angular boulders become rounded by attrition to form beaches of cobbles and pebbles. In places there are beach conglomerates consisting of basalt pebbles cemented by carbonates precipitated from seepage out of the overlying dune calcarenite. The shore platform east of Fingal Beach retains a veneer of planed-off dune calcarenite, through which underlying basalt protrudes as lava ledges and boulders. The dune calcarenite platform becomes patchy and the shore platforms flanking Rowley Cove are cut in basalt, littered with blocks and boulders. Behind Rowley Cove the steep bluff passes into a scrubby slumped area fronting an upper cliff. A gullied slope marks the zone where Selwyn Fault crosses the coast (>Fig. 21.6.5.8). From Rowley Cove the coast swings southward in high slope-over-wall cliffs along the western side of Cape Schanck. The top of the basalt rises dramatically about 40 m on the upthrow side of the fault, where the cliff becomes vertical, exposing several thick layers of basaltic lava. The steep convex upper slope of about 30° is in uplifted Pleistocene dune calcarenite, and has a cover of wind-pruned scrub.

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

21 .6 .5

⊡⊡ Fig. 21.6.5.8 Selwyn Fault (F) at Rowley Cove. (Courtesy Geostudies.)

Cape Schanck is a bold headland about 80 m high, from which lava cliffs protrude southward in an undulating promontory with platforms and ledges of basalt and bordering coves. There are also some residual cappings of dune calcarenite. Marine dissection has proceeded along joints and fissures, forming clefts and caves and isolating stacks such as Pulpit Rock from the main cliff. The Older Volcanics include massive lava flows, in places with columnar structure, interbedded with reddish horizons of weathered basalt, which formed during intervening episodes of quiescence. The basaltic land surface was then weathered, and blanketed with deposits produced when the vents reopened explosively, throwing out  masses of rock fragments, which were deposited nearby as gravel (agglomerate), and fine dust or ash which settled more widely as tuff. These pyroclastic deposits were then buried by outpourings of basalt which formed the next capping flow: in places the tuffs have been baked brick red. A magnificent section of cliffed coast cut in Older Volcanics runs NE from Cape Schanck to Flinders. The precipitous black cliffs are fronted by shore platforms, the outlines of which are strongly influenced by lava flow structures. Successive dune formations have spilled over the Cape Schanck ridge, and down the slopes bordering the incised valleys of Burrabong Creek and Main Creek. Burrabong Creek descends a deep valley in dune calcarenite, and has cut a cliff in the underlying soft freshwater limestone of early Pleistocene age, which must have formed in a lake or

lagoon behind a calcarenite barrier since removed. Seepage of carbonate-rich groundwater of from the overlying dune calcarenite has produced deposits of grey travertine over the basalt shore. Basalt pebbles have been cemented into beach conglomerate by deposition and precipitation of carbonates from this seeping groundwater. The dune capping terminates on the western slopes of Main Creek valley in a steep (about 32°) vegetated slope. Sand has spilled down into Main Creek, and has been carried downstream to Bushrangers Bay, to form a beach backed by grassy dunes. This interrupts the basalt cliffs and shore platforms, and at its eastern end is Elephant Island, a high stack of basalt with a small capping of weathered clay. Caves and arches have been cut into the cliffs, and the shore platforms are dissected by deep clefts and pools. Cliff ledges are related to lava flows. Towards Flinders the cliffs pass into vegetated bluffs, interrupted by rocky promontories developed on hexagonal columnar basalts which form ledges running out across the shore platform. At The Blowhole a deep cleft in the basalt causes waves to break, sending up clouds of spray. The bluffs are fronted by shore platforms up to 100 m wide, which include flat areas where the rock outcrops have been reduced by repeated wetting and drying (water-layer weathering). The platforms are best developed on the pink and brown tuffs, but stages in their evolution can be seen where basalt outcrops are being lowered by this weathering. There are higher ledges, especially on thicker lava horizons and some residual stacks (>Fig. 21.6.5.9).

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21.6.5

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

⊡⊡ Fig. 21.6.5.9 A stack in slightly folded columnar basalt rising from a shore platform near Flinders. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.5.10 A ledge of columnar basalt exposed on the shore west of Flinders. Blocks dislodged from the columns have been rolled and rounded as waves moved them on to the cobble beach at the mouth of Double Creek in the background. (Courtesy Geostudies.)

The shore platforms include areas flattened by waterlayer levelling. The platforms are dissected by clefts and scoured potholes. Grapeweed (Hormosira) grows on the flatter areas, and various seaweeds, including kelp, in pools and at the outer edge of the shore platform. At the mouth of the Double Creek valley there is a beach of boulders, cobbles and pebbles derived from the angular blocks that disintegrate from the hexagonal columnar basalt outcrops in

the cliffs and on the shore platform. As the basalt boulders drifted along the shore they were rounded and reduced by mutual abrasion, then piled up by storm waves to form a high beach impeding the stream outlet (>Fig. 21.6.5.10). The cliffs decline to grassy bluffs behind Ocean Beach, west of Flinders, where a beach of sand and black pebbles is backed by dunes. An outlying segment of shore platform is attached to the mainland by a gravelly swashway, ­forming

Victoria: The Nepean Ocean Coast (Point Nepean to West Head)

an intertidal tombolo shaped by the convergence of waves around the platform when it emerges at low tide. The coastal plateau runs out as a narrow promontory on West Head, the southern coast of which has a slopeover-wall profile, with slopes on deeply-weathered clays derived from the Older Volcanics and cliffs cut in hard, stratified lava and slightly deformed columnar basalt. One cove has a black beach on which the basaltic boulders released from the cliffs and shore platforms have been gradually reduced in calibre by attrition, through cobbles

21 .6 .5

and pebbles to sand. A prominent ledge of hard grey lava forms the base of West Head, which marks the beginning of Westernport Bay.

References Bird ECF (1982) Foundations. In: Hollinshed CN, Bird ECF, Goss N (eds) Lime, Land, Leisure. Flinders Shire Council, pp 1–24 Hills ES (1971) A study of cliffy coastal profiles based on examples in Victoria, Australia. Z Geomorphol 15:137–180

1357

21.6.6 V  ictoria: Westernport Bay, with French Island and Phillip Island

1.  Introduction Westernport Bay occupies part of a tectonic depression, the Western Port sunkland, between the Mornington Peninsula to the west and the South Gippsland Highlands to the east (>Fig. 21.6.6.1). On the western side the depression is bounded by the Tyabb Fault, and to the east by the Heath Hill and Bass Faults. It is more complex than the Port Phillip Bay sunkland, for it includes French Island and Phillip Island as relatively high areas and extends northward beneath the Koo-wee-rup plains as an area of swamp now largely reclaimed (Jenkin 1962). In Eocene times the Older Volcanics were deposited across the southern part of what is now Westernport Bay, and when the volcanic activity came to an end in the Oli­ gocene a wide basalt plateau extended from the Mornington Peninsula across to Phillip Island and the southern parts of French Island. Broad valleys were then cut down into

this plateau by rivers draining from north to south, and tectonic movements created the lowland within which – as in the Port Phillip area-there are Miocene and Pliocene marine deposits indicating successive incursions of the sea. In the Pleistocene there were further tectonic movements as well as sea level oscillations, so that Westernport Bay existed as a marine inlet during high sea level phases and as a coastal lowland when sea level fell. It attained its present outlines during the Late Quaternary marine transgression, about 6,000 years ago, after the sea invaded the lowlands around French Island and Phillip Island. Although tectonic movements played an important part in the evolution of Westernport Bay there is no simple relationship between the fault pattern and the present coastline. The outlines produced by tectonic deformation have been modified by subaerial processes, including ­erosion and deposition by runoff and rivers, as well as the effects of marine submergence and the shaping of coastal

⊡⊡ Fig. 21.6.6.1 The Westernport sunkland. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.6, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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21.6.6 

Victoria: Westernport Bay, with French Island and Phillip Island

and sea floor morphology by waves and currents. There was, for example, an episode in Pleistocene times when quartzose sand dunes formed in a zone that runs ­south-east from Melbourne through Cranbourne to Lang Lang and on to the northern parts of French Island. The climate was evidently then more arid, vegetation was sparse, and wind-blown sand was derived from weathered exposures of Tertiary sandstones, notably the Red Bluff Sand and the Baxter Sandstone. This sand was built into elongated ridges and parabolic dunes trending southeastward across what is now the northern part of Westernport Bay to French Island: they formed during the Late Pleistocene low sea level phase when the bay floor was dry land. When the rising sea returned to Westernport Bay some of these dunes were overwashed and rearranged into shoals. Rivers draining into Westernport Bay include small streams such as Merricks Creek, flowing from the eastern slopes of the Mornington Peninsula, those that descend from the hill country north and east of the Koo-wee-rup plains, notably the Bunyip and the Lang Lang (which formerly discharged into extensive freshwater swamps and are now conveyed to Westernport Bay by artificial channels), and the Bass River entering the south-eastern corner. These streams deliver only small quantities of fresh water and sediment to Westernport Bay, which is (like Port Phillip Bay) essentially a marine rather than estuarine inlet. Yet there are vast quantities of muddy sediment in and around Westernport Bay, forming intertidal shoals and marshlands and underlying the swampy Koo-wee-rup plains. With a more intricate configuration than Port Phillip Bay, where the intertidal zone is generally sandy or rocky, Westernport Bay has been an environment in which muddy sediment could accumulate during Pleistocene and Holocene times. The mud (consisting of silt, clay and organic matter) was derived partly from inwashed river sediment, and partly from the re-working by waves and tidal currents of fine-grained material derived from outcrops around and beneath the bay. A notable source was the clayey mantle on the weathered surface of the Older Volcanics, which are now exposed in cliffs, shore platforms and reefs around the southern parts of the bay. Onshore the basalt carries a capping of weathered clay, and a similar capping must have existed on outcrops that were submerged during the Late Quaternary marine transgression. Clay washed in from this weathered mantle was carried into the upper reaches of the bay by waves and incoming tides. Shelly organisms live in these mud deposits, and there are shell beds beneath the mud off the north and west coasts of Westernport Bay.

The shells have been sorted by waves and currents and banked alongside tidal creeks, or washed onshore to form shelly beaches, as on the north-east coast between Yallock Creek and the mouth of Lang Lang River. During the Late Pleistocene low sea level phase Westernport Bay became a lowland bordered by hill country, with French Island a central and Phillip Island a southern upland. Bass River incised its channel across this lowland to flow out of what is now the eastern entrance at San Remo, and the small streams from the Mornington Peninsula extended their courses across the emerged sea floor to meet those draining the western parts of French Island and Phillip Island in a valley that ran out beneath the wider western entrance, and across the exposed plains that floor Bass Strait. As the Late Quaternary marine transgression brought the sea back into the Westernport sunkland, waves swept sand in from the sea floor through the Western Entrance to nourish beaches along the Mornington Peninsula and Phillip Island coasts (>Fig. 21.6.6.2). To the north and east there was quieter submergence of alluvial lowlands and bordering freshwater swamps. A strait formed north of Phillip Island, and marine inlets developed and widened on either side of French Island until the two invading arms of the sea met in the north-east. Apart from a few cliffed sectors the present coast of Westernport Bay consists mainly of sandy beaches, spits and salt marshes with a seaward fringe of mangroves, fronted by extensive intertidal mudflats. The mangroves are submerged at high tide, but at low tide the sea retreats to expose a broad zone of mudflats with a patchy seagrass cover. At low tide the sea subsides into two dendritic channel systems, diverging from a tidal divide north-east of French Island (>Fig. 21.6.6.3). The floor of Westernport Bay thus has a largely tide-dominated morphology, shaped and maintained by the ebb and flow of the tides, the spring tide range increasing from 1.5 m at Flinders to over 3.3 m along the northern shore at Tooradin, but wave action has also influenced the intertidal mudflats because waves move sediment to and fro across them. In the northern parts of the bay the salt marsh and mangrove fringe has built a terrace upward and outward in front of an early Holocene coastline that was generally sandy, with some cliffed sectors. This former coastline developed about 6,000 years ago, as the Late Quaternary marine transgression came to an end. A sandy beach can be traced at the inner edge of the salt marshes along the north-western shores of the bay, around Quail Island and Chinaman Island, and on the northern shores of French Island.

 21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.2 Sand movement into the SW of Westernport Bay during Holocene times. Sand washed in from the sea floor has drifted along the coast from Flinders to Point Leo, Somers and Sandy Point, and some has been swept back on to Middle Bank by ebb tides. Sand has also drifted in to Cat Bay and along the north coast of Phillip Island, past Red Rock and Cowes to Observation Point. (Courtesy Geostudies.)

2.  Westernport Bay Coastline The coast between Flinders and Somers consists of grassy and scrubby bluffs, with cliffed headlands fronted by shore platforms, all cut into Older Volcanics, mainly lava, with some tuff and agglomerate. The basalt has generally been strongly weathered to a friable brown rock or clayey materia, but locally there are cliffs with ledges and rocky outcrops of darker unweathered lava, as on West Head at Flinders. Sandy beaches, backed by low grassy dunes, occupy intervening embayments, narrowing as they extend past rocky headlands. The beaches are 50–70 m wide at low tide, and consist of fine to medium quartzose sand with some pebbles of basalt derived from the shore platform and dune calcarenite washed in from the sea floor. Low grassy terraces, consisting of sandy material over sand and gravel beach deposits can be seen in Kennon Cove, at Shoreham Beach, and between Point Leo and Merricks Beach. The shore platforms between Flinders and Somers are subhorizontal and largely submerged at high tide and exposed as the tide falls. They have been cut partly in black

basalt and partly in brown or grey tuff. There are gaps or lower sectors in the shore platforms with beaches at Shoreham, Point Leo, Merricks and Balnarring. North of Shoreham a south-facing sandy beach extends from The Pines in a gentle curve, shaped by slightly refracted ocean swell, to Point Leo. The bluffs re-emerge as rocky cliffs on the headland at Point Leo, where there is a wide shore platform, and the beaches to the north have been shaped by waves that at high tide move in obliquely across the shore platform and drift sand northward. Between Point Leo and Balnarring there are several sectors with a backshore terrace bearing low grassy dunes that were built up by sand accretion after a phase of erosion in the 1950s, when several storms accompanied high spring tides. East from Balnarring the beach becomes a dunecapped spit that has deflected the mouth of Merricks Creek eastward (>Fig. 21.6.6.4). The position of the outlet from Merricks Creek has varied considerably. Air photographs taken in the 1920s show the mouth of the creek well to the east of its present position, opening through the beach in front of Somers. In the 1960s the present outlet was stabilised by wooden training walls, after which the

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21.6.6 

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.3 Extensive mudflats and sandy shoals are exposed at low tide in Westernport Bay, when creeks diverge from a tidal divide that runs NE from Palmer Point and flow into the deeper North Arm and East Arm channels. (Courtesy Geostudies.)

beach prograded to form a slight foreland, with an intertidal sand delta. There are nearshore reefs of weathered basalt. After dry weather the outflow diminishes and sand builds up across the creek mouth, forming a lobe directed upstream. The beach at Somers is sandy, and runs behind a wide shore platform. It is backed by steep scrub-covered bluffs, cut in Older Volcanics which pass inland as the shore platform comes to an end. The beach contin­­ues,  backed by dunes, for another 5 km to Sandy Point, off which sand

bars and shoals, partly exposed at low tide, run out SW to Middle Bank. Extensive pro­gradation in Holocene times has built a broad fore­­land with numerous beach ridges, but in recent decades there has been rapid erosion along this coast (>Fig. 21.6.6.5). The beach east of Somers is drift-dominated, with waves arriving obliquely to the shore and edging sand along the coast (>Fig. 21.6.6.2). Eastward drifting from Somers to Sandy Point takes the form of migration of a series of sandy lobes on a sinuous shore. The lobes may

 21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.4 Balnarring Beach, which has prograded in the past few years, with a grassy backshore terrace behind a convex beach profile with a wave-built berm (b). (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.6.5 Erosion of the beach east of Somers has resulted in the cliffing of backshore dunes and a low, gently concave beach profile. (Courtesy Geostudies.)

have originated from the intermittent passage of drift­­ ing sand past the mouth of Merricks Creek when it was sealed off by sand accretion, with interruptions when the creek mouth was open. Migration of sand lobes along the shore has resulted in phases of accretion as they ar­­ rived and erosion as they moved on. This drifting sand eventually reaches Sandy Point, where there has been ­sub­­stantial progradation in recent years (>Fig. 21.6.6.6) (Bird 1985).

North of Sandy Point undulating low sand cliffs mark the truncated eastern shore of the beach ridge foreland, fronted by salt marsh with a seaward fringe of mangroves that curves round into Hanns Inlet. To the north the ­mangrove-fringed salt marshes resume in front of bluffs of Baxter Sandstone, but towards Stony Point the mangroves were interrupted when they were cleared in the mid-­ nineteenth century to give access for boats that landed and took off cattle (Bird and Barson 1975).

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21.6.6 

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.6 Accretion has formed a wide beach at Sandy Point, Westernport Bay. (Courtesy Geostudies.)

North of Stony Point the coast consists of a succession of bays with mangrove-fringed salt marsh in front of wooded bluffs and low headlands, such as Crib Point, where Baxter Sandstone and ferruginous gravels are exposed. Offshore is Sandstone (Koolamadoo) Island, which consists of an anticlinal ridge in Silurian rocks, exposed in planed-off strata around the island shores and backed by grassy bluffs that were cliffs when sea level was slightly higher. At Hastings the mangroves and salt marshes have been reclaimed for port development on Long Island Point. At Yaringa a boat harbour was dredged through the salt marsh and mangroves, and out across tidal mudflats in 1967. On the NW coast similar mangrove-fringed salt marshes are advancing into Watson Bay, and they extend round Quail Island and Chinaman Island. The mangroves and salt marshes extend behind Blind Bight to Sawtell Inlet, a mangrove-fringed tidal estuary at Tooradin. The hinterland here was the extensive Koowee-rup Swamp, which was an area of reeds, rushes, paperbark scrub and shallow meres until it was drained during the later decades of the nineteenth century. South-east of the mouth of Lang Lang River the coastline consists of embayed cliffs up to 2 m high, cut in black swamp clays, generally steep or vertical, and fronted by a wave-cut ramp in firm black clay which passes seaward beneath soft mud. When these clays dry, they shrink and crack, and the polygonal fragments litter the shore as angular fragments which become rounded by wave action

to clay cobbles on the shore (>Fig. 21.6.6.7). The scalloped outlines of the Lang Lang clay cliffs may have originated before the swamp was drained, when floodwaters used to pour out as waterfalls along the shore. Offshore the dark grey mudflats, with patches of seagrass, widen on either side of a tidal divide, exposed at low spring tide, with channels diverging to the west and SE (3). At high spring tides the whole area is submerged by up to 2 m. In recent years the heads of channels on the western side have been enlarging and extending, while those to the SE have been shrinking and shortening, so that the tidal divide has been shifting SE. Red Bluff at Lang Lang is a cliff cut in Baxter Sandstone, which here consists of stratified sandstone, sandy clays and gravels. Erosion of the softer layers has led to undercutting and collapse of sandstone blocks as the cliff retreats. Waves have carried sand derived from the cliff both northward to Lang Lang jetty and beyond, and southward to nourish the growth of the recurved spit at Stockyard Point. This has grown southward by the successive addition of sandy beach and dune ridges, colonised by grasses and scrub. To the south, a peninsula of Baxter Sandstone, underlain by Older Volcanics, runs out to Corinella. At Set­ tlement Point the cliffed headland exposes the Older Volcanics, lava and tuff weathered to a friable red and ­yellow clay, which may be one of the sources of muddy sediment in Westernport Bay. The cliff is fronted by a shore platform on less weathered, or perhaps re-indurated, ­volcanic rock (>Fig. 21.6.6.8), with knobs of relatively

 21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.7 Receding clay cliffs on the coast at Lang Lang, fronted by a sloping intertidal wave-cut platform which is littered with rounded cobbles and boulders of black clay. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.6.8 Cliff and shore platform cut in weathered basalt at Corinella on the eastern shore of Westernport Bay. The weathered basalt is soft and friable in the receding cliffs, but has been superficially indurated on the shore platform. (Courtesy Geostudies.)

­ nweathered rock on small salients at the cliff base. On the u headland an emerged shell bed is one of several indications of a higher Holocene sea level in Westernport Bay, overlain by an aboriginal kitchen midden. Pelican Island is a heap of basalt blocks and cobbles exposed during the ensuing emergence, and now covered with blackwood scrub. Bays backed by low cliffs and bluffs cut in weathered basalt continue southward past Cobb Bluff to Stony Point, where a dissected lava flow runs out as an intertidal ridge

to basaltic Reef Island. Bass River flows out from a terraced valley through a wide triangular area of salt marsh interspersed with sandy beach ridges behind a shore fringed by intermittent mangroves. Within the salt marsh a cut-off meander is occupied by Spartina grass (>Fig. 21.6.6.9), introduced here from England in the early 1930s. Spartina was introduced to several Victorian estuaries: it generally failed, but has persisted in sites bordering Westernport Bay and Andersons Inlet.

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21.6.6 

Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.9 Spartina anglica terrace bordering the channel of Bass River, Westernport Bay. (Courtesy Geostudies.)

Along the southern side of the Bass valley a fault-line scarp marks the edge of the uplifted Cretaceous rocks of the South Gippsland Highlands, where bluffs and low cliffs of mudstone and Older Volcanics run out along the San Remo Peninsula, beside the eastern entrance to West­ ernport Bay.

3.  French Island French Island, with an area of about 167 sq km, consists of a broad ridge of Cretaceous rock running from SW to NE, bordered and overlain by Tertiary sands and clays, with a southern fringe of Older Volcanics, and extensive areas of Pleistocene quartzose dunes (Jenkin 1962). Tortoise Head, at the SW corner, is a former island of Older Volcanic rocks, mainly basalt, attached by a beach of sand and gravel, bordered on the eastern side by salt marshes with a mangrove fringe. The headland is cliffed on the western and southern shores, which are exposed to attenuated ocean swell as well as SW storm waves, but these pass into grassy bluffs on the more sheltered eastern and northern flanks. The bluffs were formerly cliffs on Tortoise Head, shaped by stronger wave action when sea level was higher. A landslide in weathered basaltic clays on the western side has pushed out a festoon of rocky boulders across the shore. To the north there are bluffs and a shore platform cut in weathered Older Volcanics at Tankerton, but beyond the pier the bluff recedes behind a low coastal plain ­bearing sandy beach ridges, fringed by remnants of an eroded salt

marsh, a line of old mangroves marking their former seaward limit (Bird and Barson 1975). Mangroves and salt marshes resume along the north shore of French Island to Palmer Point, where they are up to a kilometre wide. At Palmer Point the salt marshes still show relics of rectangular enclosures built in the nineteenth century as evaporation ponds for salt production, but they were soon abandoned. East of Palmer Point low beach-fringed cliffs and bluffs pass behind a cuspate sandy foreland at Spit Point (also known as Sandy Point), where the northern shore has been truncated by wave action. The SE coast of French Island has cliffed sectors at Blue Gum Point and Stockyard Point, and a sandy beach with occasional mangroves behind soft mudflats. On the embayed south coast there are cliffs and platforms cut in Cretaceous mudstone from Stockyard Point past Red Bluff almost to Long Point, and on the northern point of Elizabeth Island. Long Point and Peck Point are headlands with cliffs and reefs of Older Volcanics, which also dominate outlying Elizabeth Island, and form the causeway that runs out to Rams Island. Between the headlands are wide curving sandy bays, but the shore is too exposed for salt marshes or mangroves, except in the lee of Tortoise Head.

4.  Phillip Island Phillip Island lies across the southern part of Westernport Bay, between the wide Western Passage and the narrower Eastern Entrance. Its coast consists largely of cliffs and

 21.6.6

Victoria: Westernport Bay, with French Island and Phillip Island

bluffs bordering a gently undulating plateau that rises southward to elevations of 40–50 m along the bold Bass Strait coast. The island is dominated by Older Volcanic rocks (Eocene). Successive lava flows spread across the landscape of Phillip Island, and in cliff sections, especially along the south coast, they are seen as horizontal or gently dipping (rarely more than 5°) layers of solid or cindery grey to black rock, with intervening faintly stratified horizons of red, brown or grey tuff and agglomerate formed by the deposition of volcanic ash and gravel. On the south coast of the island cliffs and steep bluffs have been cut into basaltic lava, while the tuffs form intervening gentler slopes. At Point Grant solid lava flows can be seen in the cliffs and in The Nobbies, a chain of steep-sided islets that rise above the wide bouldery intertidal shore platform. There is a small capping of Pleistocene dune calcarenite. Pink and grey tuffs and agglomerates underlie a dissected lava flow; their colours are due to varying degrees of oxidation of iron compounds contained in the volcanic rocks. Offshore, remnants of basaltic lava flows form Seal Rocks. The south coast of Phillip Island has cliffy headlands and grassy bluffs on Older Volcanics (>Fig. 21.6.6.10). Shore platforms are strongly influenced by geological structure, many segments coinciding with the upper surface of a lava flow. Some are storm wave platforms, shaped by strong wave attack which has dislodged and removed weathered ⊡⊡ Fig. 21.6.6.10 Bluffs on the south-east coast of Phillip Island were cliffs during a mid-Holocene higher sea level phase, behind segments of slightly emerged shore platform still visible at high tide. (Courtesy Geostudies.)

rock material and cut out cliff-face ledges along flat or gently dipping lava surfaces, as on Helen Head. To the east Woolamai Surf Beach is backed by Pleis­ tocene dune calcarenites and capping Holocene dunes of the Woolamai isthmus, up to 100 m high, which attaches the granitic upland of Cape Woolamai to the mainland. In contrast with the extrusive volcanic rocks, the pink granite of Cape Woolamai was intruded into Palaeozoic rock formations, probably during the Devonian. The granite is a coarse crystalline rock, with numerous vertical and horizontal joints which intersect the crystalline rock at various angles, and have been etched out by weathering to produce cracks and crevices between blocks, columns and pinnacles of granite. The granite cliffs of Cape Woolamai are higher (about 75 m) and more rugged on the south-west side, which faces the strongest waves, than on the less exposed southeastern and eastern shores. Marine erosion has penetrated along joint planes, forming alternating gorges and buttresses. Storm-tossed boulders are found along the cliff tops, proof of the fury of wave attack during storms. Rock stacks formed of pillars and columns of granite occur along the east coast of Cape Woolamai, separated from the cliff by marine erosion along joint planes (>Fig. 21.6.6.11). The Cape Woolamai granite ends northward on the shores of Cleeland Bight. Dunes have spilled across the Woolamai isthmus to the shore of Cleeland Bight, supplying sand to form wide beaches backed by grassy low foredunes. Longshore drifting carried sand northward as far as the road bridge at Newhaven.

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Victoria: Westernport Bay, with French Island and Phillip Island

⊡⊡ Fig. 21.6.6.11 The granite cliffs on the east coast of Cape Woolamai plunge below low tide level, and have no beach fringe. (Courtesy Geostudies.)

On the west coast of Phillip Island Grossard Point is a rocky promontory formed where hard, dark, little-­ weathered basalt outcrops at the cliff base and on the adjacent platform, but the cliff face is receding in soft, weathered tuff which has decomposed to form cracking clays. Outcrops of red tuff form small cliffs and shore platforms between Penguin Rocks and Cowes, where three small headlands interrupt the sandy beach. Sandy beaches are almost continuous, interrupted only by minor rocky promontories, from Cat Bay to Ventnor, and on past Cowes and Erehwon Point to form the northern shore of a compound recurved spit which culminates in Observation Point (>Fig. 21.6.6.2). East of Cowes the nearshore waters become very shallow, the intertidal zone broadening into Cowes Bank, a sandy area about 5 km long and up to 450 m wide. It is dominated by large, asymmetrical sand bars and intervening troughs, some of which contain muddy sediment up to 5 cm thick. On the shores of Rhyll Inlet, in the lee of the Observation Point spit, mangroves and salt marshes are drained by tidal creeks which converge eastwards into mudflats exposed at low tide. Between Rhyll and Newhaven the east coast of Phillip Island is sheltered from strong wave action, and has

­ angrove-fringed salt marshes behind wide intertidal m sand and mudflats in Swan Bay, Denne Bight and Reid Bight, backed by bluffs which sweep out to successive partly cliffed promontories of weathered basalt and tuff at Chambers Point, Pleasant Point, and Long Point. At low tide extensive mudflats and sandflats are exposed, with creeks draining to the deeper channels. The bluffs continue round to Churchill Island, an embayed hilly island of weathered basalt and tuff, separated from Phillip Island by a shallow strait where the basalt is overlain by mudflats. To the south is a bay with mangroves and salt marsh, backed by low bluffs which run out to form the coast past Newhaven to the Phillip Island bridge.

References Bird ECF (1985) Recent changes on the Somers-Sandy Point coastline. Proc R Soc Vic 97:115–128 Bird ECF, Barson MM (1975) Shoreline changes in Westernport Bay. Proc R Soc Vic 87:15–28 Jenkin JJ (1962) The geology and hydrogeology of the Westernport area. Mines Department, Underground Water Investigation Report No. 5

21.6.7 South Gippsland, Wilsons Promontory and Corner Inlet

1.  Introduction The South Gippsland coast consists of cliffed capes and wide asymmetrical embayments running south-east to the granitic uplands of Wilsons Promontory, backed by the Corner Inlet depression. The hinterland is dominated by the Cretaceous sandstones and mudstones of the South Gippsland Highlands (also known as the Strzelecki Rang­ ­es), geologically similar to those of the Otways. They outcrop in deeply dissected hill country formed by subaerial denudation of a series of uplifted domes and fault-bounded blocks, with intervening areas of subsidence in which Tertiary and Quaternary sediment has accumulated. The Tertiary rocks include areas of Older Volcanics, but these formations have had little effect on the coastline.

2.  South Gippsland Coastline In the western part of the South Gippsland Highlands the Bass Block, bounded by the Bass and Kongwak Faults,

forms an upland that runs SW to reach the coast in the cliffs between San Remo and Kilcunda. The cliffs and shore platforms are cut in Cretaceous sandstones and mudstones (>Fig. 21.6.7.1) similar to those of the > Otway coast, but backed by an undulating plateau incised by valleys. The coastline is crenulate’, with bold anticlinal headlands and valley-mouth coves. Cliff recession has truncated some of the valleys, so that streams pour out as waterfalls, and caves and natural arches have been cut out along joint planes. Within the small coves are locally derived pebble and cobble beaches, but Shelly Beach occupies a larger cove where white shells derived from nearshore reefs have been heaped upon the shore. The Punchbowl is a steep-sided conical hollow about 50 m across and 30 m deep, formed as a result of subsidence of the roof of a cave, in the sandstone cliffs: it is still slowly sinking. A col across a headland marks a valley dismembered by cliff recession, the former course of the stream incised behind the following embayment. Below the Coalmine Valley is a cove with a natural arch and a pedestal remnant of an earlier arch rising from the shore platform.

⊡⊡ Fig. 21.6.7.1 The cliffed coast east of San Remo, cut into Cretaceous sandstones and mudstones, with sandy beaches in coves and shore platforms on headlands. The road leads to the Punchbowl. (Courtesy Geostudies.)

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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South Gippsland, Wilsons Promontory and Corner Inlet

At Kilcunda the cliffs decline, the shore platform becomes intermittent and beyond the small valley of Bourne Creek the beaches lengthen, backed by grassy foredunes. Powlett River drains a wide fault-bounded lowland and reaches the sea by way of a gap through the dunes. The beach then continues, backed by Holocene dunes disrupted by blowouts spilling inland, exposing outcrops of Pleistocene dune calcarenite. Shore platforms of Cretaceous sandstone and mudstone, with some coal seams, reappear off Coal Point, and extend in front of the cliffed coast to Cape Paterson. Generally the headlands coincide with basal outcrops of harder and more massive sandstone, which also form higher segments of shore platform, while the mudstones have been excavated to form embayments behind which sandy beaches have accumulated, often backed by grassy dunes. Cape Paterson is a low promontory bordered by cliff and shore platforms cut in layered sandstones, interrupted by sandy surf beaches backed by high scrubby dunes that have advanced across it. The cliffed Bunurong coast to the NE has bold headlands where streams flow out over waterfalls and sandy beaches are backed by steep grassy bluffs in the coves. Eagles Nest is a large rock stack rising from the shore platform off a bold cliffed promontory (> Fig. 21.6.7.2). To the east successive declining cliffy promontories stand behind Flat Rocks, a wide, boulder-strewn shore platform

cut across dipping Cretaceous sandstones and mudstone. Fossil remains of dinosaurs and early mammals have been found on this coast. The cliffs pass into grassy bluffs, relics of a Pleistocene coastline that extends inland behind the Inverloch coastal lowland. Eastward drifting beach sand has deflected the mouth of Wreck Creek, and Point Norman is a lobate foreland bearing parallel dune ridges. Inverloch, bordering the entrance to Andersons Inlet, has a sandy beach, along which successive lobes of sand have drifted in recent decades. Andersons Inlet is the most clearly defined estuary on the Victorian coast, with extensive tidal mudflats (16 km2) exposed at low tide on either side of a meandering channel that runs from the reed-fringed river at Tarwin Lower to the sea at Inverloch. A broad lagoon forms as the tide rises, with a 38-km shoreline that is partly sandy and partly swampy. A sandy threshold has been washed into the entrance by waves and inflowing tides. There have been changes in the outline of Point Smythe, across from Inverloch, where a sandy foreland forms intermittently at the NW end of the dune-capped barrier separating Andersons Inlet from Venus Bay. In the 1960s Point Smythe was trimmed back by waves and current scour, and ended in a receding dune cliff, but in the subsequent decade a large sandy foreland grew out in front of the cliffed dunes, with grasses and shrubs colonising

⊡⊡ Fig. 21.6.7.2 Eagles Nest, a stack rising from the shore platform on Cretaceous sandstone on the Bunurong Coast. (Courtesy Geostudies.)

South Gippsland, Wilsons Promontory and Corner Inlet

wave-built berms to build up a series of hummocky dunes. This persisted until 2003, when the sandy foreland was again removed by erosion. The changes are related to fluctuations in the pattern of channels at the entrance to Andersons Inlet. Along the northern shore of Andersons Inlet, Townsend Bluff is a cliff, fronted by a sloping intertidal shore platform cut across Cretaceous sandstones by wave abrasion. This truncates a volcanic plug, 20 m in diameter, intruded into the Cretaceous sandstones. Spring tide range at Inverloch is about 1.7 m, and tidal movements and salinity decline into Andersons Inlet. Sandy sediment near the entrance gives place to mudflats produced by deposition of silt and clay from Tarwin River, and threaded by winding tidal channels. Reedswamp is extensive around the river mouth, but mangroves grow on the shores of the estuary, fronted by mudflats with a patchy cover of seagrasses, which are otherwise remarkably sparse in Andersons Inlet. Rice grass (Spartina anglica) was introduced from Britain in 1962 and soon spread rapidly (>Fig. 21.6.7.3). It began to build the mudflats into extensive bordering terraces close to high tide level, covered by a grassy sward (Bird and Boston 1968). A long surf beach runs SE along the shores of Venus Bay, backed by high grassy Holocene dunes over Pleistocene dune calcarenites which emerge on the shore towards the southern end to form high cliffs fronted by shore platforms,

⊡⊡ Fig. 21.6.7.3 Early stages of Spartina anglica invasion, Andersons Inlet. Brick dust has been placed on the mud to measure subsequent accretion. (Courtesy Geostudies.)

21.6.7

from which rises two rugged stacks called Arch Rock as the name implies, these were formerly linked as a natural arch that has collapsed. Steep bluffs pass behind Morgan Beach, at the southern end of which begin the rugged cliffs and rocky shores bordering the Hoddle Range, terminating in Cape Liptrap. The high cliffs at Cape Liptrap are cut in strongly folded hard Lower Devonian sandstones and mudstones, the Liptrap Formation. The harder elements persist as ribs, and minor anticlines and synclines can be seen in varying stages of dissection on the shore. Cliffs cut in steeply ­dipping strata have basal abrasion notches and the less resistant outcrops have been worn down to form shore platforms (> Fig. 21.6.7.4). There are steep beaches of locally derived grey cobbles and pebbles in the coves, and below the western side of the Cape are remnants of emerged shore platform, probably of Pleistocene age, 3–4 m higher than the present shore platform. East of Cape Liptrap, the cliffs become steep vegetated bluffs, mantled by earthy rubble formed by weathering, and fronted by a shore platform which has been cut across almost vertical hard sandstone and mudstone strata, which break up into angular gravel. Along the coast the shore platform becomes very irregular, with ribs of harder rock in the tightly folded strata, deep transverse clefts, rocky coves with cobble beaches, and caves. Towards Grinder Point the Cambrian greenstones (metamorphic rocks consisting of ancient lavas, tuffs, agglomerates and shales

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21.6.7

South Gippsland, Wilsons Promontory and Corner Inlet

⊡⊡ Fig. 21.6.7.4 Shore platform cut across Devonian sandstones and shales at Cape Liptrap. (Courtesy Geostudies.)

much altered by intense heat and pressure) emerge and are overlain by stratified Waratah Limestone, of Devonian age. The bluffs then decline to a gentler slope, and incised streams drain to the shore along Maitland Beach, a steep beach of storm-piled cobbles, which is partly vegetated and could have originally formed when the sea stood at a higher level. The bluffs steepen into cliffs towards Bell Point, where the Waratah Limestone is overlain by the darker Bell Point Limestone. Both formations have been corroded by sea spray, and cut back to form flat shore platforms, with residual stacks that show basal notches and visors produced by solution processes (> Fig. 21.6.7.5). The Devonian limestone is flanked by Ordovician sandstones and mudstones along the shore northward past Digger Island, where they include an outcrop of older limestone, to Walkerville. Bird Rock is one of a group of spiky islands and stacks of Waratah Limestone running out from the headland. At Walkerville there are overgrown coastal quarries in Waratah Limestone. The sea has cut back into coastal ledges behind sandy coves, and locally the overlying earthy weathered material has slumped to the shore. The limestone comes to an end at The Bluff, where it is faulted against steeply dipping mudstones of the Liptrap Formation, planed off in the cliff section and overlain by earthy rubble that collapses on to the beach. There are outcrops of yellow calcareous sand over grey quartzose sand indicating that dunes spilled across the Cape Liptrap ridge in a similar way to those at Cape Otway and Cape Paterson.

To the north the scrubby bluffs behind Waratah Bay were cliffs when the sea stood at a higher level. They are fronted by a narrow sandy beach and wide intertidal shore platforms cut in the Liptrap Formation. The curving bluffs have been shaped by south-westerly waves refracted round the Cape Liptrap promontory to approach the coast from the SE. The bordering sandy beach continues to the SE along the spit that borders Shallow Inlet. At Sandy Point the Waratah Bay beach is backed by a series of low grassy dune ridges and swales with amplitudes of up to 1.5 m. The mouth of Shallow Inlet has migrated to and fro in an area of sandy banks and shoals. Swell moving into this part of the shore of Waratah Bay is refracted over a protruding sandy tidal delta. In contrast with Andersons Inlet, Shallow Inlet receives only minor stream inflow, and is essentially a marine rather than an estuarine system. Spring tide range is about 2 m, and at high spring tides the water covers about 18 km2, while at low spring tides there is water only in a meandering channel covering only about 3 km2. Tidal ventilation is thus of the order of 32 million m3 and the ebb and flow of this volume of water generates strong currents. At high tide the lagoon is broad enough for waves to be generated across it, especially by the prevailing westerly winds, and these have trimmed back low cliffs on parts of the eastern shore. The strip of land between Waratah Bay and Corner Inlet is the Yanakie isthmus, capped by hummocky dunes, partly covered by scrub and woodland, with many swampy hollows. There is a large unvegetated dune area which was

South Gippsland, Wilsons Promontory and Corner Inlet

21.6.7

⊡⊡ Fig. 21.6.7.5 Mushroom Rock showing notch and visor cut in Devonian limestone near Bell Point, Walkerville. (Courtesy Geostudies.)

probably the outcome of such a sudden accession of sand to the eastern shore Shallow Inlet a century or two ago. As on the other Victorian promontories, older quartzose dunes emerge from beneath the dune calcarenites to form heathy dune topography on the hillside south of Darby River. The underlying granites appear on Shellback Island, and emerge, capped by dune calcarenite, in Tongue Point to the south. This marks the beginning of Wilsons Promontory.

3.  Wilsons Promontory The Devonian granites of Wilsons Promontory formed as a batholith, an intrusion of magma (molten rock) about 390 million years ago into Lower Palaeozoic rocks which have since been removed. The magma cooled slowly and crystallised as granite, a mixture of quartz, felspars, micas and accessory minerals. It was exposed as the overlying rock formations were worn away, then deeply dissected by runoff and rivers to form mountainous ranges with peaks more than 700 m high, ridges that run out to coastal headlands, and incised valleys opening to sandy embayments. The coastal landforms of Wilsons Promontory are influenced by the pattern of fractures and joints in the granite, which have guided dissection, both by stream incision and marine erosion, to shape bays and headlands, coves, clefts and caves. Where the joints are widely spaced the massive granite outcrops as cliffs and ledges. Blocks of granite form tors and buttresses where weathering has

decomposed the rock along intervening joints and the weathered material has been washed away. Stripping of weathered material continues on hillside slopes, which have aprons of gritty colluvial downwash consisting of weathered material containing angular grains of quartz and felspar (Hill and Joyce 1995). On the west coast of Wilsons Promontory the beach and dunes in Leonard Bay consist almost entirely of quartz, with small proportions of other minerals derived from granite. The beach is known as Squeaky Beach because when it is dry it emits a musical sound when walked on. It is a very well sorted medium sand with rounded quartz sands, which under a hand lens look like tiny glass marbles. The quartz grains have become well rounded as the result of intense and prolonged wearing by wave action of material derived from the granitic weathered mantle that extended below present sea level. As this was submerged and reworked by the waves, the less resistant feldspar and associated minerals were destroyed. The dunes behind Squeaky Beach are of similar white quartz sand, blown up from the beach, while on the slopes to the south there are old quartzose dunes above a coast that is now rocky, even at low tide. The other beaches on the south-west coast of Wilsons Promontory are quartzose, but have varying proportions of calcareous sand washed in from the sea floor. It is possible that the rocky offshore islands of the Glennies Group have prevented this sand reaching Leonard Bay, by impeding drifting by the predominant south-westerly ocean swell.

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21.6.7

South Gippsland, Wilsons Promontory and Corner Inlet

In Norman Bay (> Fig. 21.6.7.6) the beach is wide and very gently shelving, consisting of sands that are partly calcareous (up to 40% carbonate), backed by parallel Holocene foredunes, grassy at first, then covered by scrub and woodland, mainly coast tea tree (Leptospermum laevigatum). From the top of Mount Oberon these parallel ridges can be seen to be backed by older, probably Pleistocene, parabolic dunes bearing heath vegetation, and sheets of quartzose sand blown some way up the hillside, and bordered by Tidal River. Little Oberon Bay is a sandy cove of coarse white sand, forming a steep beach. It is quartzose, with some felspar and flecks of black tourmaline, together with other heavy minerals derived from the weathered mantle of the granite, washed down from the slopes of Mount Oberon and exposed in the cliff behind the cove. Oberon Bay has more calcareous beach sands, and the backing dunes include dissected ridges of Pleistocene dune calcarenite, exposed in major blowouts that run inland to spilling noses of sand that are burying older vegetated dunes. The steep-sided granite peninsulas that run out between these embayments have rocky and bouldery shores. Erosion has been guided by joint planes in the granite, and there has been exfoliation (the peeling off of successive surface rock layers) on the exposed granitic slopes. Beneath the lighthouse an exfoliated slope of bare granite descends steeply, plunging to deep water close inshore. There has been almost no marine erosion of this outcrop in the 6,000 years of Holocene stillstand. Similar features are seen on outlying islands and stacks of granite

in Bass Strait, notably the steep plunging cliffs around Skull Island and the huge pillar of Rodondo, rising 350 m out of Bass Strait. The east coast of Wilsons Promontory is dominated by steep rocky slopes, with several promontories. Valleys descend to Sealers Cove and Waterloo Bay, which have beaches of quartzose sand shaped by ocean swell moving in from the SE; the calcareous sand seen on the western coast evidently came in from the west across the floor of Bass Strait, and was not received on this east-facing shore. The beaches are backed by thickly vegetated swamps, fed by streams from the high granite slopes. To the north, 5-Mile Beach fringes a barrier with low parallel beach ridges and foredunes, then an extensive swamp on the site of a former lagoon, within which is a high wooded knoll, a former island. The NE corner of Wilsons Promontory consists of a lowland with numerous parallel foredunes marking intermittent progradation of a sandy shore, and carrying scrub, woodland and heath vegetation on soils that show stages in podzolic evolution. At Entrance Point the foredunes have been truncated on the north-facing shore, which passes into granite below the slopes of Mount Singapore.

4.  Corner Inlet Corner Inlet, like Port Phillip Bay and Westernport Bay, occupies a tectonic depression bounded by faults to the ⊡⊡ Fig. 21.6.7.6 Norman Bay, Wilsons Promontory, showing the wide surf beach between granite spurs, backed by parallel dune ridges beneath tea tree woodland. (Courtesy Neville Rosengren.)

South Gippsland, Wilsons Promontory and Corner Inlet

north and west, with Yanakie isthmus attaching Wilsons Promontory to the south. It is essentially a shallow submerged plain, an infilled basin or alluvial lowland. As the tide falls, extensive sand and mudflats are exposed and a dendritic system of deeper channels converges eastward to the entrance. Many of these channels originate on tidal sandflats and mudflats near the coast, but Franklin River drains into one of them. There is little evidence of rapid siltation, but rivers certainly carried sediment into Corner Inlet during 1934 and 1952 floods. Several granite islands rise through the shoals; Doughboy Island has a small shelly beach, Granite Island a guano-coated northern shore, and Bennison Island an emerged rock bench. The eastern shores of Corner Inlet have sandy beaches below the granitic slopes, and between rocky headlands are sandy barriers fronting extensive freshwater swamps in valley mouths and embayments that were formerly lagoons. Chinamans Beach has sands that vary in colour from yellow to reddish brown, and probably include material derived from erosion of submerged podzolised sandy deposits, notably humate horizons enriched by iron oxides washed down from the leached topsoil. Off Millers Landing nearshore granitic islets with tors and boulders show granular disintegration by recurrent wetting and drying, and associated salt plucking, which eventually planes them off. To the SE the shore is fringed by rushy salt marshes, which pass southward behind stunted mangroves (Avicennia marina) that reach their southernmost global limit here at 38° 54' 24'' S.

⊡⊡ Fig. 21.6.7.7 Dune sand spilling on to the southern shore of Corner Inlet. (Courtesy Geostudies.)

21.6.7

In the southern part of Corner Inlet some of the intertidal shoals bear salt marsh and a few mangroves. On the SW shore the mangroves give place to sandy beaches bordering the dunes that have spilled across the Yanakie isthmus (> Fig. 21.6.7.7). To the north soft sandy mudflats front a swampy area bearing paper-bark woodland. Backshore erosion has cut microcliffs behind sloping abrasion ramps in peaty clay, and there is a veneer of sand. This becomes coarser northward to Townsend Point, where weathered granite reappears in low cliffs and wave-cut ramps. The coarse sand, derived from weathering granite, forms a beach and swash bar, and is locally cemented to beach rock by carbonates precipitated from a former shelly beach capping. The beach curves out to a low cuspate foreland at Duck Point, on which very low wave-built beach ridges with an amplitude of a few centimetres are marked by vegetation zones. These are truncated on the northern shore, and there is convergent drifting out to the arrow-shaped point. To the west of Duck Point is a cliff cut into cemented Tertiary quartz gravels. At Red Bluff, near Yanakie, a north-facing cliff shows the Wilsons Promontory granite strongly weathered to a sandy clay (quartz sand grains in a matrix of clay formed by the decomposition of felspars) and stained red by iron oxides, passing beneath steeply dipping weathered red and yellow Palaeozoic (probably Ordovician) shales and sandstones. Offshore there are grey granite tors surrounded by soft mudflats. A long

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21.6.7

South Gippsland, Wilsons Promontory and Corner Inlet

⊡⊡ Fig. 21.6.7.8 Mangrove island in the north of Corner Inlet, showing mangroves spreading into a central salt marsh. (Courtesy Geostudies.)

sandy beach backed by scrub-covered bluffs ends in an arrow spit at Swan Point. The west coast of Corner Inlet is generally muddy and marshy, with some mangroves, backed by low-lying terrain with swamp paper-bark woodland, and embanked areas of land reclaimed for agriculture. Patches of Spartina grass have spread from plantings made on the shores of Tidal Creek, south of Foster, in 1930. The NW coast of Corner Inlet consists of salt marsh in front of a low depositional terrace, which has the appearance of an emerged tidal flat, formed during a Holocene phase of higher sea level, and backed by bluffs marking the mid-Holocene coastline. At Foster Beach this low terrace still has a cover of swamp paper-bark, and the seaward margin has been cut back by wave erosion. Behind the midHolocene bluff the land rises to a broad Pleistocene terrace, into which Bennison Creek, Franklin River, Toora Creek and Agnes River have each incised valleys that open to mangrove-fringed tidal inlets. Evidence of possible subsidence has come from studies of the patterns of mangrove and salt marsh vegetation on the low islands in Corner Inlet. These islands evolved from shoals where accretion of sand and mud was aided by colonising vegetation, initially

mangroves and then salt marsh. However, on several of the islands the outer edge of the mangroves is now eroding, and mangroves are spreading back into the salt marsh (> Fig. 21.6.7.8). This reversal could be an indication of a recent rise in sea level, augmented by local subsidence (Vanderzee 1988). The coastal plain widens east of Toora, and on Barry Point there are hummocky Pleistocene quartzose dune ridges interspersed with swampy corridors. This is part of an early Pleistocene barrier topography that runs eastward behind Port Welshpool towards the sandy barrier coastline of East Gippsland.

References Bird ECF, Boston KG (1968) Spartina in Victoria. Vic Nat 85:11–18 Hill SM, Joyce EB (1995) Granitic regolith and landscape evolution of Wilsons Promontory, Victoria. Pro R Soc Vic, pp 1–10 Vanderzee M (1988) Changes in salt marsh vegetation as an early indication of sea level rise. In: Pearman G (ed) Greenhouse, planning for  climate change. CSIRO Division of Atmospheric Research, Melbourne, pp 147–160

21.6.8 Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

1.  Introduction East from Corner Inlet the escarpment bounding the South Gippsland Highlands, on Cretaceous rocks along the Gelliondale Monocline, recedes behind a widening coastal plain of Pliocene and Pleistocene sands and clays. This plain is crossed by Tarra River, Bruthen Creek, Merrimans Creek and some smaller streams, and shows scarps and undulations that may be the result of Pleistocene and Holocene tectonic movements. They include the Lake Wellington depression to the NE, part of the down-warped zone occupied by the Gippsland Lakes. Between Yarram and Woodside, and NE to Sale and Bairnsdale, the landscape carries scattered ridges of Pleistocene quartzose dune sand, often of parabolic form, which may have been coastal sand barriers rearranged by westerly winds during a more arid period (Jenkin 1968). Large quantities of sand have been deposited on the East Gippsland coast as beaches, barriers and dunes in Quaternary times. The sand is predominantly quartzose, with only small proportions of calcareous material, and as there are only very limited cliffed sectors, and the rivers drain mainly into coastal lagoons (except for the Snowy, which delivers some sediment to the coast), the sand must have come from the floor of Bass Strait. As on other sectors of the Victorian coast, shoreward movement of sand has occurred as the result of onshore wind action during low sea level phases, and wave deposition during and after successive marine transgressions. There is a contrast between these quartzose sands, derived ultimately from weathered granitic formations, and the calcareous sands on the Victorian coast west of Wilsons Promontory. The East Gippsland coastal barriers consist of successively formed dune ridges parallel to the coastline, except where they have been interrupted by the development of blowouts, which have in places grown into parabolic dunes or coalesced in transgressive sand sheets. Some of the dunes are active and mobile, but most are held beneath a cover of vegetation. Dunes that formed during the Pleistocene have

deep podzolic soil profiles with humate horizons, and generally carry heathy vegetation, while the Holocene dunes have much shallower soil development, and are held in place by grasses and shrubs, scrub and woodland.

2.  The East Gippsland Coast East of Corner Inlet is a series of barrier islands, each consisting of numerous sub-parallel beach ridges and foredunes arranged in complex, often intersecting patterns, interspersed with swampy corridors and fringed on the landward side by estuarine channels with areas of salt marsh and mangroves (>Fig. 21.6.8.1). The ridges mark stages in the progradation of a sandy shore; their truncation indicates episodes of erosion and re-shaping. There are similar sandy deposits on the coastal plain from Port Welshpool to Port Albert, Manns Beach and St. Margaret Island with deeply leached podzolic profiles and a cover of heathy woodland. They are interrupted by the estuaries of the Albert and Tarra Rivers, which incised their courses across them during the Late Pleistocene low sea level phase. They are therefore of Pleistocene age. The barrier islands that developed in front of this Pleistocene coast are largely of Holocene age, consisting of multiple beach and dune ridges built within the past 6,000 years, together with salt marshes and mangrove swamps in areas of muddy accretion sheltered from strong wave action. East of Port Welshpool a tidal divide exposed during the ebb was formerly used to drive cattle to and fro from Snake Island. This is the largest of the sandy barrier islands, with dune ridges similar to those seen on the NE corner of Wilsons Promontory. Townsend Point on the southern shore is a lobate foreland consisting of young parallel foredunes, and remnants of earlier such forelands can be found within the pattern of ridges on Snake Island, several of them truncated before later ridges were appended. The present beach widens along the shore east of Townsend Point, and there has been progradation here since nautical charts were made in the nineteenth century.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.8, © Springer Science+Business Media B.V. 2010 (Dordrecht)

21  .  6  . 8

⊡⊡ Fig. 21.6.8.1 Coastal landforms east of Corner Inlet. (Courtesy Geostudies.)

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Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

Sunday Island is similar to Snake Island. The pattern of sand ridges shows that it developed in several stages, with episodes of truncation by laterally migrating tidal channels: the youngest part is a narrow, segmented Holocene barrier island known as The Drum, off the southern shore. There is a soil and vegetation sequence across the parallel dune ridges confirming their increasing age to landward (Turner et al. 1962). Spring tide range is generally less than 1.5 m along the  East Gippsland coast but it increases to nearly 3 m at Port Albert, and the ebb and flow currents strengthen accordingly. The channels between the barrier islands are  maintained by tidal scour, and the openings to the sea  are the outcome of a contest between wave action, tending to build up the sandy barrier, and the transverse  tidal currents. Port Albert Entrance now opens between paired sand spits to a broad shallow sea area from  which sand is still being washed shoreward by swell that arrives from the south. Progradation has formed a very wide sandy shore (>Fig. 21.6.8.2). To the east the outer barrier is interrupted by gaps at Kate Kearney Entrance, then Shoal Entrance, both tidal inlets of variable form and  position, with seaward intertidal and subtidal sand deltas. The SW end of the Ninety Mile Beach shows parallel ridges that curve landward (Jenkin 1968), as the beach diverges into a long, narrow sand spit on the eastern side of  Shoal Inlet. Meandering of the tidal channel south of St.  Margaret Island cut through these ridges from the

⊡⊡ Fig. 21.6.8.2 Prograded sandy coastline south of Port Albert where sand has been washed in from nearshore shoals. (Courtesy Geostudies.)

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l­andward side and in 1961 a new entrance formed at McLaughlins Inlet. This breaching shortened the Ninety Mile Beach from 89.2 miles (143.5 km) in 1960 to 85 miles (136.8 km). This entrance has been migrating NE, with receding cliffs cut across the foredunes on the northern side and sand deposition on the southern side. Bruthen Creek flows into an estuarine lagoon fringed by salt marshes and scattered mangroves behind the sandy outer barrier to the east. Mangroves are not found east of here; the various lagoons and inlets have only salt marshes, often dominated by rushes (Juncus spp.) (>Fig. 21.6.8.3). The outer barrier sands are largely quartzose, with generally less than 10% of carbonate sand. In the absence of a sandy supply from eroding cliffs or river discharge, the sand has come from the sea floor. It has been derived largely from the quartzose weathered mantle of the granite outcrops around the Bass Strait islands to the SW. Shoreward drifting of sand by wave action during and after the Late Quaternary marine transgression resulted in the accre­­ tion  of the Ninety Mile Beach backed by successivelyformed parallel foredunes, along a prograding coastline (>Fig. 21.6.8.3A, B). However, erosion is now dominant, and the coastline is receding as sand is lost offshore and alongshore (>Fig. 21.6.8.3C, D). Beach erosion may have been accentuated by tectonic subsidence east of Corner Inlet, and by sea floor and coastal subsidence resulting from the extraction of Bass Strait oil and natural gas. At McLaughlins Beach a meandering tidal channel has hl cut into the landward margin of the innermost (and thus

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Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

⊡⊡ Fig. 21.6.8.3 Beach and dune evolution, Ninety Mile Beach. Sand moved shoreward from sea floor shoals as sea level rose late in Pleistocene times, forming a beach and backing dunes (a, b). When sea level stabilised (stillstand) shoreward drifting ceased (c), and as the concave profile shaped by wave action extended shoreward the beach and dunes began to erode (d). (Courtesy Geostudies.)

oldest) sand ridge of the outer barrier to expose basal gravelly material and some calcarenite. This shows that the first deposits laid down as the Late Quaternary marine transgression came to an end were coarser and more calcareous than those added during subsequent sandy progradation. Similar features have been noted at other points along the inner shore of the outer barrier. The Ninety Mile Beach is typically 20–30 m wide at low tide, facing SE, with one or more intermittent ­nearshore

sand bars. It is a sandy surf beach with patches of shelly gravel, and has been shaped largely by south-easterly ocean swell. Nevertheless, it shows alternations of longshore drifting, to the north-east when waves are generated by southerly winds, and to the south-west when the waves come in from the east. There are variations in the width and profile of the beach, which widens round sand lobes that form and migrate generally NE. Behind such a lobe the beach profile is convex and the backshore dunes have a vegetated seaward slope, sometimes with a grassy terrace

Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

or incipient foredunes. Between the lobes, where the beach is narrower and steeper, with a concave profile, the backshore dunes are generally cut back as a cliff of crumbling sand. In one of these sectors, NE of Jack Smiths Lake, the backshore dune cliff shows dune sand underlain by a horizon of black swamp clay that formed in a lagoon or swamp to landward, and has been exposed by coastline recession. Although the dunes are generally cliffed, segments of new grassy foredune develop in summer, sometimes persisting for several years until the next major storm removes them. Between McLaughlins Inlet and Seaspray the coastal foredunes are backed by a low corridor of lagoons and swamps, with a bluff on the landward side marking a former (Late Pleistocene) coastline. Jack Smiths Lake may occupy an area of tectonic subsidence, and after wet weather raises the water level, overflow cuts a temporary outlet through the Ninety Mile Beach. Lake Denison, to the north-east is smaller, but similar. Merrimans Creek comes down to the coast at Seaspray, and has an outlet through the Ninety Mile Beach that varies in width and depth with the strength of fluvial outflow, being occasionally sealed off in droughts. Beyond Letts Beach the coast curves gradually round to a more southerly aspect, and eastward drifting by SW waves then predominates. The beach also shows seasonal cut-andfill alternations, the profile being built up by constructive swell in summer to become convex, often with a berm of firmly packed sand (>Fig. 21.6.8.4), then lowered and cut

⊡⊡ Fig. 21.6.8.4 During calm weather gentle ocean swell builds a berm of firm, fine sand on the Ninety Mile Beach, firm enough to land a light aircraft. (Courtesy Geostudies.)

21.   6  . 8

back during storms, especially in the winter, to form a backshore cliff and a gently inclined, sometimes concave slope (>Fig. 21.6.8.5), with sand removed to nearshore bars. The Holocene outer barrier shows lateral variations in width and the number and size of parallel foredunes: between Woodside and Seaspray there are only one or two ridges, and in places a single ridge has been transgressive, driven inland over peaty deposits that formed in a backing lagoon and swamp, and are now exposed along the Ninety Mile Beach (>Fig. 21.6.8.6). Near Letts Beach the parallel dune ridges increase to a maximum of 13, but between Ocean Grange and Lakes Entrance there are again only one or two. Typically the foredunes rise 5–10 m above the intervening swales, and are spaced at intervals of 20–50 m. Behind the outer barrier is a corridor of lagoons and swamps, backed by a Late Pleistocene coastline (>Fig.  21.6.8.6,  B) that consisted of a bluff (formerly a cliff) truncating Pliocene anhld Pleistocene formations (Jenkin 1968). Merrimans Creek flows to the coast by way of a valley incised into a Pleistocene terrace, emerging from the bluff to flow through the lagoon to the sea by way of a small gap through the outer barrier dunes. The Ninety Mile Beach curves east towards Lakes l Entrance, where stone breakwaters border an artificial channel cut through the outer barrier in 1899 to provide ship access to and from the Gippsland Lakes (> Victoria: The Gippsland Lakes ). At the same time a looped sand bar formed outside the entrance, and has persisted. Sand has

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Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

⊡⊡ Fig. 21.6.8.5 Backshore cliff cut on Ninety Mile Beach during a storm. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.8.6 At Seaspray the Ninety Mile Beach is backed by a single grassy foredune, and then a swampy lowland overlooked by a bluff B – marking the Late Pleistocene coastline. (Courtesy Geostudies.)

accumulated on either side of the Lakes Entrance breakwaters (>Fig. 21.6.8.7), forming triangular areas on which foredunes have developed since 1899 (Bird 1978). The Pleistocene bluff which lies behind the township of Lakes Entrance curves out to the cliffed sandstone headland at Red Bluff. At high tide this marks the eastern limit of the Ninety Mile Beach, but at low tide the sandy beach continues past it, and similar beaches, backed by dunes, resume to the east as far as Cape Conran. There is

an overall eastward drifting of sand along the East Gippsland coast to Cape Howe. Lake Tyers is a branched lagoon formed by submergence of two incised valleys to form Toorloo and Nowa Nowa Arms, and enclosed by the outer barrier. The mouth is often sealed off by a sand bank, and the lake then becomes almost fresh, but from time to time an outlet is cut by a stormy sea or overflowing floodwaters from the rivers, or by excavation of a trench by local people.

Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

To the east the outer barrier is backed by a long swampy corridor, Ewings Morass, fed by several small streams, some of which have built deltas. This was formerly a lagoon, similar to Lake Reeve, and extending from Lake Tyers to Lake Corringle, but it has been filled by sedimentation and swamp encroachment. It is backed by a bluff which runs in behind Lake Corringle to merge with the valley-side slopes overlooking the Snowy delta. North of Orbost, the Snowy River emerges from a deep and narrow valley to flow across a broad deltaic plain, with steep natural levees bordering the sandfloored  river channel and extensive backswamp depressions backed by bluffs cut into gravel-capped terraces on Miocene limestones and sandstones. The river reaches the sea by way of a migratory outlet through the sandy outer barrier. The Snowy River carries large quantities of sand and some gravel downstream during floods, when sand is swept into the sea. Some of this coarse angular sand has subsequently been washed up on to nearby beaches. The sandy beach resumes beyond Point Ricardo, but is soon interrupted by the larger granitic Cape Conran, which has outlines influenced by jointing. Dock Inlet is a freshwater lake, generally about 2 m deep, where a former coastal embayment has been completely cut off from the sea by the dune-capped outer barrier. Pearl Point is a dune-capped promontory fringed by Ordovician slates and sandstones which interrupt the sandy beach. To the east the dune-capped sandy barrier resumes, and extends in front of Sydenham Inlet, a lagoon

⊡⊡ Fig. 21.6.8.7 Sand accretion at Lakes Entrance. (Courtesy Geostudies.)

21.   6  . 8

up to 2.5 m deep, fed by Bemm River. The lagoon shores are partly sandy and partly reed and rush-fringed. Tamboon Inlet is another estuarine lagoon, up to 6 m deep, with rocky shores as well as beaches of grey sand brought down by Cann River, which has built a small reedfringed delta. The eastern shore abuts a high granite ridge. To the east the ocean beach is interrupted by Clinton Rocks, granite outcrops beneath a bold 12 m cliff cut in bedded Pleistocene quartzose dune sands with firmly cemented humate horizons; it resumes, backed by extensive parabolic dunes which widen inland, as the coast curves out towards Point Hicks. Formerly known as Cape Everard, Point Hicks is a broad promontory of granite with a wide rocky and bouldery shore, which includes some Ordovician metamorphic outcrops west of the lighthouse. Petrel Point has cliffs cut in dark metamorphosed Ordovician sediment and outcrops of granite. Here, as elsewhere along the East Gippsland coast the granite forms rugged, bouldery shores rather than shore platforms. Rame Head is, like Point Hicks, a high granite promontory capped by scrub-­ covered dunes. The coastline swings north to the mouth of  Wingan Inlet, another estuarine lagoon with steep rocky  shores behind a dune-capped Holocene sand barrier. Wingan River enters the Inlet after descending over rapids across the granite outcrop, and has built a small delta into the head of the lagoon, while towards the outlet is a threshold (tidal delta) of inwashed marine sand. The outlet is narrow but persistent, impacted against a granite promontory.

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Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe)

⊡⊡ Fig. 21.6.8.8 Corridor planation on mudstone outcrops in the Ordovician rocks near Little Rame Head. (Courtesy Geostudies.)

Between Little Rame Head, where there are outcrops of black quartzite, and Mallacoota, the SE facing coast has cliffs cut in strongly folded Ordovician ­sand­­stones, shales, slates and cherts. These cliffs form the ­seaward edge of a terrace 20–30  m above sea level, backed by a ­rising hinterland. Beneath the even-crested cliffs the ­foreshore on strongly folded rocks is rugged, with shore platform corridors cut in shales and slates sepa­­rated  by ribs and ridges on the harder sandstones and quartzites (>Fig. 21.6.8.8). Rock outcrops in cliffs show honey­ combing and irregular weathering due to wetting and ­drying and salt spray, and ironstone joint fillings stand out as an intricate fretwork near Seal Cove and Shipwreck Cove. There are boulder and cobble beaches, and abrasion potholes scoured in the shore platforms. The cliffed coast has only small patches of sandy beach, although sand is moving along the sea floor past Little Rame Head towards Mallacoota. At Bastion Point the cliffs turns north, passing into bluffs on which the seaside resort of Mallacoota stands. Mallacoota Inlet is an extensive branched estuarine lagoon system fed by the Genoa and Wallagaraugh Rivers. The dune-covered spit on the eastern side of the entrance formed during the Holocene, and there is a tidal delta of inwashed sand surmounted by marshy islets just inside the entrance, which is variable and obstructed by a sand spit and bars. East of Mallacoota a narrow barrier with grassy ­foredunes widens into a broad sandy fringe with cuspate

forelands in the lee of Tullaberga and Gabo Islands, shaped by the refraction of a southerly ocean swell. The grassy foredunes are interrupted by blowouts that become larger eastwards, growing into low, then high parabolic dunes, and eventually an extensive eastward-drifting sand sheet, with sets of spilling transverse ridges (Rosengren 1981). The dune fringe is backed by rush-fringed freshwater lakes at Lake Barracoota and Lake Wau Wauka, where a stream has been ponded back by the arrival of mobile dunes across a valley to form an elongated lake. Offshore, Tullaberga Island is fringed by rocky shores with outcrops of Ordovician rocks and red Devonian granite, and Gabo Island is a larger, partly dune-covered granite island, seaward of a marine channel: Nineteenth century maps showed it as a tombolo. Dunes extend along the coast to Cape Howe and are spilling on top the shore in New South Wales.

References Bird ECF (1978) The geomorphology of the Gippsland Lakes Region. Ministry for Conservation, Environ­mental Studies Series. Publication 186, Melbourne Jenkin JJ (1968) The geomorphology and upper cainozoic geology of southeast Gippsland, Victoria. Geological Society of Victoria, Memoir 27 Rosengren NJ (1981) Dune systems on cuspate forelands, East Gippsland. Proc R Soc Vic 92:137–47 Turner JS, Carr SGM and Bird ECF (1962) The dune succession at Corner Inlet. Proc R Soc Vic 75:17–33

21.6.9 Victoria: The Gippsland Lakes

1.  Introduction The Gippsland Lakes are a large and complex group of coastal lagoons, separated from the sea by broad sandy barriers. They include Lake Wellington (138 sq km; shoreline length 60 km), Lake Victoria (110 sq km; shoreline length just over 100 km), Lake King (92 sq km; shoreline length 160 km), and a number of smaller lagoons associated with  extensive swamps on a low-lying depositional coastal plain (Bird 1978). These lakes are fed by five main rivers, the Latrobe and the Avon flowing into Lake Wellington, the Mitchell, Nicholson, and Tambo into Lake King; and together these drain a catchment of about 20,600 sq km, extending into the Eastern Highlands of Victoria (>Fig. 21.6.9.1). Much of the catchment above the 150 m contour is still forested, but the lower country is mainly grazed or cultivated (irrigated) farmland, with some residual patches of bush and heathland. Here the river valleys are incised into a coastal plateau which borders the southern slopes of the Eastern Highlands, and ends in a marginal bluff overlooking the Gippsland Lakes. The valley floors are low-lying and swampy, with natural levees bordering river channels and deltas protruding into the lakes. The coastal barriers are dune-covered, and bordered on the seaward side by the Ninety Mile Beach (> Victoria: The East Gippsland Coast). The lakes are generally shallow, much of Lake Wellington being less than 4 m in depth, and much of Lake King less than 6 m. Lake Victoria occupies an elongated furrow up to 10 m deep that extends eastwards across the southern part of Lake King, past Metung, and through to Lakes Entrance. In McLennan Strait, which links Lake Wellington to Lake Victoria, the depth attains a maximum of 11 m as the channel curves past Seacombe. The deepest point in the Gippsland Lakes is 15 m, in a scour-hole east of Shaving Point at Metung. The coastal barriers (>Fig. 21.6.9.2) include an inner and an outer barrier between the lakes and the sea, and remnants of a prior barrier, so called because it was initiated before the inner and outer barriers developed to enclose the Gippsland Lakes in their present form. Each barrier is surmounted by dune topography, largely fixed in position by a cover of  scrub, woodland and heath

vegetation. Each has a ­ground-pattern of dune ridges running parallel to the curved outline of the Ninety Mile Beach, but interrupted by blowouts and parabolic dunes. The crests of the dunes are generally less than 10 m above mean sea level, but locally they are much higher, and some of the summits exceed 30 m. The coastal barriers represent the accumulation of a large amount of sand on the East Gippsland coast in Late Pleistocene times: it has been suggested that a tsunami may have contributed to this accumulation (Bryant and Price 1997). When the explorer Angus McMillan arrived in this area the Gippsland Lakes were linked to the sea only by way of an intermittent and variable outlet through the eastern end of the barrier, between Cunninghame (now the town of Lakes Entrance) and Red Bluff. River floods used to build up the level of the lakes until water spilt out across the low-lying section of the outer barrier east of Cunninghame, opening a natural outlet. In dry weather reduced river flow and high evaporation lowered the level of the lakes, and this outlet, no longer maintained by outflow, became sealed by sand deposition. This impeded navigation, and so an artificial channel was cut through a narrow section of the outer barrier opposite Jemmys Point, near Cunninghame, to provide a navigable artificial entrance to the lakes (>Fig. 21.6.9.3). After this opened in 1889 the old natural outlet soon became sealed off permanently by sand deposition. Cun­ninghame Arm, alongside the town of Lakes Entrance, is the shrunken remnant of a former navigable channel that once led to the natural outlet (Bird and Lennon 1989). The climate of Gippsland is mild and maritime. Meteorological records from Bairnsdale, a station typical of the Gippsland Lakes region, show a range of mean temperature from 8.3°C in winter (July) to 19.6°C in summer (February), and a mean annual rainfall of 669 mm, evenly distributed through the year. East Sale has a mean annual rainfall of 602 mm, and Lakes Entrance 736 mm. Rainfall is in fact more variable than the mean figures suggest, for there are occasional droughts, usually in late summer and autumn. Prevailing winds are westerly, but easterly winds blow mainly during the summer months. The hydrology of the Gippsland Lakes is influenced by local climate, the regime of the inflowing rivers, and the movement of water in and out of the artificial entrance.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.6.9, © Springer Science+Business Media B.V. 2010 (Dordrecht)

⊡⊡ Fig. 21.6.9.1 East Gippsland, showing the catchment of the Gippsland Lakes. (Courtesy Geostudies.)

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21.6.9   Victoria: The Gippsland Lakes

Victoria: The Gippsland Lakes  

21.6.9

⊡⊡ Fig. 21.6.9.2 The coastal barriers in the Gippsland Lakes region, with a cross-section (AB) showing the relationship between the Late Pleistocene prior and inner barriers and the Holocene outer barrier, resting upon a Pleistocene land surface. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.9.3 The artificial entrance to the Gippsland Lakes, looking south over Bullock Island, an artificial island formed by dumping sand dredged from the looped bar outside the entrance. (Courtesy Geostudies.)

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21.6.9  

Victoria: The Gippsland Lakes

Water is supplied to the lakes by rain and rivers and removed by evaporation; it flows in and out of the artificial entrance according to the relative levels of lake and sea. Water level fluctuations are complicated by seasonal variations. Fluvial discharge reaches a maximum in late winter and early spring, when runoff is augmented by the melting of snow in the Eastern Highlands. River flooding occurs most frequently at this time, and may temporarily raise the level of the Gippsland Lakes by up to 2 m, inundating extensive areas of the bordering lowlands. After major floods the lake levels take several days to subside. Rainfall shows little variation from month to month, but in winter evaporation is much reduced. Consequently, lake levels rise during the winter months, and outflow at the artificial entrance is strengthened by discharge of this additional water. During the summer the lake levels fall, their minimum level being determined by sea level at the artificial entrance. Tides in the Gippsland Lakes depend on the flow of water in and out of the artificial entrance. At the Lakes Entrance breakwaters the tides are diurnal and similar to those along the Bass Strait coast, tide range increasing from 0.6 m at neap tides to about 1.0 m at spring tides. Tidal movements are not transmitted far into the lakes. In calm, dry weather the surface of the lakes stabilises close to mean sea level, and measurements then indicate a tide range of about 10 cm at Metung, with diurnal tides are barely perceptible farther west, in Lake Victoria and Lake Wellington. After heavy rain and river flooding raise the level of the lakes, water pours out to sea through the artificial entrance even when the tide is rising. Strong winds and variations in barometric pressure also modify lake levels, wind action moving water across the lakes to build up their level to leeward and lower it to windward by as much as 60 cm above and below calm-weather lake level. When the wind drops the calm-weather level is restored after a succession of standing wave oscillations (seiches) which cause irregular changes of water level around the lake shore, and sharp alternations of current flow through McLennan Strait. These variations, and the currents that accompany them, are not related to the astronomically-generated tides observed in the adjacent ocean, although they mingle with tidal movements transmitted through the artificial entrance and into the eastern part of Lake King. Currents generated by the ebb and flow of tides through the artificial entrance attain velocities of up to five knots (2.2 m/s) and similar velocities are generated by strong winds blowing over the lakes. River floods produce currents of up to seven knots (3.0 m/s) off river mouths, in narrow straits, and out through the artificial entrance, but these are short-lived. Considerable erosion may be accomplished by such currents: several metres of land were

scoured away from the shoreline of Lake Wellington near the mouth of the Avon by exceptional floods in 1952. The lakes are oriented in such a way that the prevailing westerly winds blow across long stretches of water, and produce waves up to a metre high, with lengths of up to 10 m. The less frequent easterly winds also generate wave action across the lakes towards east-facing sectors. Ocean swell does not penetrate into the Gippsland Lakes, except to a very limited extent, just inside the artificial entrance. Salinity in the Gippsland Lakes is largely determined by the meeting and mixing of fresh water from rain and rivers and salt water flowing in from the sea through the artificial entrance. There is a transverse salinity gradient, increasing towards the artificial entrance, with variations related to seasonal and longer-term climatic conditions. The opening of the artificial entrance in 1889 initiated a new salinity regime with unrestricted inflow of sea water in dry seasons. Salinity has increased in the Gippsland Lakes since the opening of the artificial entrance, and there have been various ecological and geomorphological changes resulting from this salinity increase. The lakes have various types of shoreline. Some sectors are fringed by reeds or rushes, others by beaches of sand or sand and shingle; some show active cliffing, others are backed by steep and relatively stable bluffs with a forest, scrub or pasture cover. Some are advancing (prograding), others are receding (retrograding) and others are stationary. In the last few decades erosion has become widespread on the shores of Lake King and Lake Victoria, and there is evidence shorelines that previously prograded are now being cut back (Bird 1983). Much of the land adjacent to the lakes is low-lying and swampy, subject to periodic flooding. There are tracts of reedswamp (reed vegetation growing in shallow water) salt marsh (low herbaceous halophytic vegetation, frequently submerged), and swamp scrub (dense scrub vegetation on land a few cm above calm weather lake level, occasionally flooded). Mangroves are not present: the narrowness of the artificial entrance has evidently prevented the arrival of mangrove seeds. Tracts of featureless low plain are found at a slightly higher level (1–2 m above calm-weather lagoon level) and consist of deposits of sand and silt, inundated only occasionally by exceptionally high river floods. There is a marginal bluff on the landward side of the Gippsland Lakes and their associated lowland. Between Metung and Lakes Entrance this bluff is a steep forested slope that was a sea cliff before the barriers were deposited in front of it (>Fig. 21.6.9.4). To the east it runs out to a promontory at Red Bluff, marking the eastern end of the Ninety Mile Beach. West of Metung the marginal bluff passes along the sides of the river valleys, where it has been

Victoria: The Gippsland Lakes  

21.6.9

⊡⊡ Fig. 21.6.9.4 The marginal bluff, formerly a sea cliff, extending west from Jemmys Point, Lakes Entrance. (Courtesy Geostudies.)

locally ­undercut by the river, as at Eagle Point Bluff, where  Pleistocene gravels cap Upper Tertiary sandstones (>Fig. 21.6.9.5). It runs behind low-lying country north of Lake Victoria and Lake Wellington, where again it was formerly a coastal cliff exposed to the waves of Bass Strait. Enclosed by coastal barriers, the Gippsland Lakes became a relatively sheltered environment, within which sediment was deposited, mainly by inflowing rivers. When the last marine transgression came to an end there were estuarine inlets at the mouths of each of these rivers, and sediment derived from their catchments was deposited in them, forming alluvial valley floors. By the time European explorers arrived in the 1840s the catchment was largely forested, and the rivers were delivering only fine sediment, mainly silt and clay, to the Gippsland Lakes. The rivers have thus reclaimed sections of their previously drowned valleys by sedimentation, producing valley floors with river channels bordered by natural levees and backswamps. The natural levees pass downstream into deltas (silt jetties) that protrude into the lakes. The Latrobe River has built a cuspate delta into the south-west of Lake Wellington, and the Avon River has formed a similar, smaller delta in the NW of that lake. Both deltas grew as the result of sedimentation in bordering reedswamp that spread forward into the lake, and was followed by swamp scrub. Deposition of silt was accompanied by the formation of organic material beneath seasonally decaying reedswamp and brushwood peat beneath the swamp scrub. In Lake King a similar delta formed at the mouth of the Tambo River (>Fig. 21.6.9.6), while the Mitchell River

⊡⊡ Fig. 21.6.9.5 Eagle Point Bluff, showing Pleistocene gravels over Pliocene sandstones). (Courtesy Geostudies.)

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Victoria: The Gippsland Lakes

⊡⊡ Fig. 21.6.9.6 The Tambo delta has lost its former reedswamp fringe and is now eroding. (Courtesy Geostudies.)

built an impressive digitate delta extending 6 km out into the lake, with silt jetties on either side of the river channel (>Fig. 21.6.9.7). The vegetation on the Mitchell River silt jetties is now mainly pastureland but there are residual patches of swamp scrub and some areas of salt marsh. Reedswamp is present sparsely alongside the river channel, with larger clones near the river mouth, but it is missing from the outer shoreline of the delta. When it was originally surveyed in 1849 the Mitchell delta was broader in outline than it is now, and much of the bordering shore was ­reedswamp-fringed. The disappearance of the reedswamp fringe has been followed by erosion of the delta shoreline (>Fig. 21.6.9.8). Comparison of the 1849 map with outlines shown on modern air photographs shows that shoreline erosion has reduced the delta area to little over half it original extent (Bird and Rosengren 1971). Some parts are very narrow. The eastern part of the delta has been dissected into a chain of small islands, and in August 1919 a gap was cut through the northern jetty opposite Eagle Point Bluff by overflowing river floods (>Fig. 21.6.9.9). This was enlarged by later floods, and outflow through this gap has deposited silt in a lobe extending out into the shallow waters of Jones Bay top the north. In the absence of reedswamp vegetation no new delta has formed here. The disappearance of shoreline reedswamp from deltas in Lake King and Lake Victoria and its reduction in Lake Wellington, are correlated with increasing salinity in the Gippsland Lakes following the opening of the artificial

entrance in 1889. Other effects of this salinity increase include the replacement of freshwater weeds, notably Val­ lisneria spiralis, by marine seagrasses such as Zostera spp., and the invasion of wave-splashed delta shores by salt marsh species. Swamp land borders the shores of the Gippsland Lakes, particularly on sectors sheltered from strong wind and wave action. It is extensive (up to 6 km wide) west and south of Lake Wellington, having prograded on shores not  exposed to waves generated by the prevailing westerly and south-westerly winds. Swamp land has also developed in sheltered inlets, along the margins of the straits and narrow lagoons, and adjacent to the river deltas. It formed as the result of an encroachment succession from reedswamp to swamp scrub and woodland around the lake shores (>Fig. 21.6.9.10). With the disappearance of shoreline reedswamp erosion has become prevalent (>Fig. 21.6.9.11). The present outlines of the Gippsland Lakes originated when the Late Quaternary marine transgression invaded low-lying areas between the dissected prior and inner barriers and the marginal bluff, and submerged parts of the valleys of inflowing rivers. With the completion of the outer barrier the configuration of the lagoons began to change as the result of wave and current action, causing erosion and deposition around the lake shores. Bays were excavated, and intervening spits grew, notably the large recurved spit that separated Lake Wellington from Lake Victoria. With sea water largely excluded the lakes became fresher, and reedswamp spread around their shores,

Victoria: The Gippsland Lakes  

21.6.9

⊡⊡ Fig. 21.6.9.7 The Mitchell River silt jetties, a digitate delta built out into Lake King. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.9.8 The eroding shoreline of the Mitchell delta in Eagle Point Bay. (Courtesy Geostudies.)

i­ nitiating the encroachment of swamp land. There was a period of shrinkage and shallowing of the Gippsland Lakes as sediment, mainly from rivers, was deposited, and the shores of deltas and swamps advanced into the lakes. After the artificial entrance was cut, and salinity increased, the shoreline vegetation was much reduced, erosion ensued, and the process of excavating embayments and building up intervening spits revived.

If land and sea levels remain in their present relative positions, continuing sedimentation will eventually fill the Gippsland Lakes, but their future mode of evolution will depend largely on ecological conditions. If they were freshened as the result of natural or artificial changes, shoreline reedswamp would revive to halt much of the erosion that is now going on and promote swamp encroachment and consequent shrinkage of the lakes. If they remained brackish,

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Victoria: The Gippsland Lakes

⊡⊡ Fig. 21.6.9.9 The Cut, a gap cut by floodwaters in the northern arm of the Mitchell delta below Eagle Point Bluff in 1919 and subsequently enlarged. Discharge through this gap has deposited silt and driftwood to form a submerged delta in Jones Bay. (Courtesy Geostudies.)

⊡⊡ Fig. 21.6.9.10 Reedswamp ­encroachment on the southern shore of Lake Wellington. (Courtesy Geostudies.)

or if their salinity increased (e.g. if shoreline erosion on the Ninety Mile Beach opens new gaps through the outer barrier), their shorelines would remain without vegetation and growth of spits between eroded bays will lead to segmentation: the conversion of the lakes into a chain of smaller and shallower, rounded lagoons. Under these conditions, sediment carried into the lakes would not be trapped in bordering reedswamp, but would settle on the

lake floor, resulting in shallowing of the lakes. Either way, the Gippsland Lakes will ultimately give place to a depositional coastal plain traversed by the engrafted estuaries of the several rivers as they flow to an outlet on the coast of Bass Strait. Alternatively, if the forecast global sea level rise proceeds, the coastal barriers will be breached and removed, and the marine embayment that existed early in Pleistocene times will be resurrected.

Victoria: The Gippsland Lakes  

21.6.9

⊡⊡ Fig. 21.6.9.11 Eroding swamp scrub land on the northern shore of Lake Victoria. (Courtesy Geostudies.)

References Bird ECF (1978) The geomorphology of the Gippsland Lakes Region. Min­ istry for Conservation, Environmental Studies Series, Publication 186 Bird ECF (1983) Shoreline changes in the Gippsland Lakes. Proc R Soc Vic 95:227–235

Bird ECF, Lennon J (1989) Making an entrance: The story of the artificial entrance to the Gippsland Lakes. James Yeates, Bairnsdale Bird ECF, Rosengren NJ (1971) The disappearing Mitchell delta. Proc R Soc Vic 84:189–197 Bryant EA, Price DM (1997) Late pleistocene marine chronology of the Gippsland Lakes region. Aust Phys Geogr 18:318–334

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21.7 Tasmania

Joanna Ellison

1.  Introduction While the main island of Tasmania lies between 39° 40ʹ and 43° 20ʹ S, this only island state of Australia extends nearly to the Victoria coastline because it includes the large Bass Strait Islands of King and Flinders, as well as 233 smaller islands (Brothers et al. 2001) and the southern outlier Macquarie Island at 54° 30' S. As an island state Tasmania has more coastline per unit land area than others, with a total coastal length of 6,473 km, excluding Macquarie Island, and no location more than 115 km from the coast. The main island coastline, some 2,236 km in length, exhibits strong contrasts in environmental geomorphology in response to wide variations in geological structure and exposure.

2.  Coastal Morphology of Tasmania Tasmania has a cool temperate climate, with prevailing winds and swell approaching from the west and southwest, which can be of very high energy owing to storm generation on the Antarctic Polar Front and uninterrupted fetch. The swell is refracted round southeastern Tasmania to arrive from the south and southeast on the east coast, and round northwestern Tasmania to arrive from the northwest on the northwest coast. However, the central inflection of the north Tasmanian coast greatly reduces longshore movement in this section, and the northeast is influenced by easterly storms that drive significant longshore drift from east to west. Mean spring tide ranges are small in southern Tasmania (0.9 m at Hobart, 0.5 m in Port Davey) and larger on the north coast (2.1 m at Waterhouse Island, 2.3 m at Georgetown at the mouth of the Tamar, 2.6 m at Devonport, and 2.5 m at Burnie).

3.  The Tasmanian Coastline The southern coast of Tasmania exposes the oldest rocks in the state: Pre-Cambrian quartzites. While the coast is

steep and rugged, sheer cliffs are rare, and small beaches occur in embayments behind headlands of boulders to sand. The most exposed island group is of pre-Cambrian quartzite formation, the Maatsuyker Group, off the southwest coast. To the east the Huon river discharges into the narrow D’Entrecasteaux Channel, which is bordered by small headlands and coves, some of which have shingle beaches derived from disintegrated dolerite. Southeastern Tasmania has long peninsulas between intricate bays with pockets of beaches. The Derwent River at the head of Storm Bay at Hobart has one of the largest discharges in Australia; the bay was initiated as a graben during the early Tertiary 65 million years ago. At present the Derwent and the Huon are large drowned river valleys (rias), marine inlets formed by the Holocene marine transgression from 13,500 to 6,000 years ago. Sea level since the mid-Holocene has been relatively stable, though this is still under discussion. Ocean swell moving from the south and east has been refracted to shape beaches that face in various directions around Frederick Henry Bay, notably the curved, sandy Seven Mile Beach, which borders a dune-capped barrier at the head of the bay. Bruny Island and Maria Island off southeastern Tasmania each consist of two large islands connected by a sand tombolo. South Bruny Island features towering dolerite cliffs, and is connected by a 7-km sand isthmus to Bruny Island to the north. On both Bruny Islands cliffs face Storm Bay to the east, which are eroded fault scarps adjacent to the depressed Storm Bay region. To the east is the Tasman Peninsula, similar to Bruny Island and presenting dolerite cliffs to the prevailing weather from the southwest. Spectacular dolerite columns are exposed along extensive cliffs, particularly at Cape Pillar (>Fig. 21.7.1) and Cape Hauy. These are Australia’s highest coastal cliffs. At sea level there is a Holocene shore platform. Permian sedimentary siltstone and mudstone are exposed in embayed sections of the Tasman peninsula, eroded by wave action to form blowholes, arches, and ­collapse indentations (>Fig. 21.7.2).

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_21.7, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 21.7.1 Columnar Jurassic dolerite at Cape Pillar.

⊡⊡ Fig. 21.7.2 Tasman Arch at Eaglehawk Neck, a collapsed sea cave more than 50 m high in siltstone.

Tasmania

The geology of Maria Island is complex, with late Devonian granite headlands on the east coast between richly fossiliferous Permian limestones and sandstones. Jurassic dolerite forms the peaks and ridges of the north island, which is connected to the south island by a 1-km sand isthmus. The east coast of Tasmania is relatively sheltered, with sand beaches that grow in width towards the north, and sand barriers. The granite Freycinet Peninsula (>Fig. 21.7.3) borders the eastern shore of Great Storm Bay, and waves from the south are refracted to form the curved Nine Mile Beach sand barrier at its head. With shack development, ⊡⊡ Fig. 21.7.3 Granite intrusions now exposed as part of the Hazards Range on the Freycinet peninsula.

⊡⊡ Fig. 21.7.4 Confined mouth of St. Georges Bay, with sand deposits blocking boat passage.

21.7

such sand barriers are thought to be one of the most vulnerable areas of Tasmania’s coast to projected sea level rise. North of Freycinet are extensive sandy beaches derived from erosion of largely granite catchments. This is the driest coastline of Tasmania, and influenced by the warm East Australian Current, particularly in summer. River mouths here demonstrate problems with sand bars blocking boat access, a problem which is also prevalent along most north Tasmanian estuaries, which could also be associated with reduced river discharge as a result of increased catchment water use and dams. At St. Georges Bay a complex late Pleistocene and Holocene barrier (>Fig. 21.7.4) shelters

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the estuarine lagoon of the George River (Davies 1985). Here is located the fishing town of St. Helens, and options for maintaining a navigable entrance to this through sand deposits in the bay mouth are still unresolved. Tin mining in the hinterland has resulted in the movement of sand and gravel waste down rivers. The Golden Fleece arm of the George River lagoon has been filled, and a cuspate delta has grown at the mouth of George River. The northeastern coastal plain of Tasmania consists of low outcrops of granite and dolerite emerging through extensive sedimentary plains formed from Tertiary gravels and Quaternary marine sands. This coastline of Tasmania is the richest in Tasmania in sand deposits as a result of the Pleistocene dune systems developed during late Quater­ nary sea level change. Bass Strait was exposed during Pleistocene sea level low stands, and arid climatic conditions during the last glacial stage caused formation of extensive dunes, which remain today on the north coast of  Tasmania. Fields of longitudinal dunes occur over 350 km2 of NE Tasmania, extending in places more than 10 km inland (>Fig. 21.7.5). These have a consistent WNW-ESE (c. 280o–100o) alignment, indicating the predominant wind direction during time of formation during the Last Glacial Stage (Bowden 1983). The sand ridges are narrow, relatively straight, up to 20 km long, 60 m wide, and 10 m high. Several large active parabolic dunes also occur in northeast Tasmania, caused by reduced plant cover with overgrazing, fire, and drought. Parabolic dunes at Water­ house Point are among the fastest moving in the world, up to 14 m/year, and some threaten to move over farmland. Flinders Island to the north is the largest of the eastern Bass Strait Islands, primarily constructed of granite

mountain peaks to the west, and sedimentary plains on the east coast, these consisting of dune and lagoon topography. The age of this landform decreases from the centre of the island towards the east. Cameron Inlet and other large lagoons (Sellars Lagoon, Logan Lagoon) occur between Pleistocene and Recent sedimentary units (Kershaw and Sutherland 1972), separated from Bass Strait by Holocene deposits of Lackrana sand, which extend along the entire east coast (>Fig. 21.7.6). The majority of Tasmanian Ramsar sites are in north and east Tasmania, and are coastal lagoons behind Holocene sand barriers such as these. Some are associated with salt marsh, which tends not to be extensive in Tasmania owing to the micro to mesotidal ranges. On the north coast of Tasmania wave energy is low, with dissipative shallow bathymetry offshore, resulting in largely spilling waves and low gradient beaches. Here Tasmania’s largest estuary is the Tamar (Edgar et al. 2000), extending 71 km inland from Bass Strait to Launceston, where it receives discharge from the South Esk and North Esk rivers, with a catchment area comprising over 20% of Tasmania (about 10,000 km2). The Tamar valley is a graben downfaulted between dolerite horsts to the west and east, to form a geologically confined estuary with a deep, well-defined channel, and tidal amplification from a mean tidal range of 2.34 m at George Town to 3.25 m at Launceston. It is bordered by tidal mudflats and wetlands. Rice grass (Spartina anglica) was introduced in the 1940s to improve navigability, and today the Tamar contains Australia’s largest rice grass infestation of 440 ha, while the second largest occupies the Rubicon estuary to the west. In the Tamar this has been recently shown to have trapped about 1,193,441 m3, with between 14–28% of that being

0

5

N

10 KM

⊡⊡ Fig. 21.7.5 Linear dunes, lunettes, and interglacial marine sands of northeastern Tasmania. (From Bowden and Colhoun 1984.)

Tasmania

21.7

⊡⊡ Fig. 21.7.6 View from south to north along the east coast of Flinders Island, showing major lagoons enclosed by Holocene sand barriers.

⊡⊡ Fig. 21.7.7 Shingle Beach at Lillico Beach, west of Devonport.

Spartina-derived organic matter, reflecting the high inorganic sediment inputs from the catchments. While nearly all Tasmanian beaches are of sand, shingle occurs on Lillico Beach west of Devonport (>Fig. 21.7.7). This section of coastline from Table Cape to Point Sorrell features deposits of shingle and sand, upon which some

coastal towns are constructed. These deposits were laid down during the Pleistocene, which included several periods of glacial formation and melting in the upper areas of the river catchments such as the Mersey and the Forth. During ice melt, fluvioglacial outwash brought shingle and sand down the catchment (Davies and

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Tasmania

Hudson 1987), and deposited these over areas at the present day coast. Wave action has since eroded this loose sedimentary coastline, releasing shingle into the beach environment. At Turners Beach, eroded dune scarps show horizontally orientated shingle strata, intermixed with sand. Turners Beach may be in an earlier phase of geomorphic evolution, of erosion of shingle rich dunes, than the all-shingle Lillico Beach to the east. The largest western Bass Strait Island is King Island, consisting of a low Plateau of pre-Cambrian quartzite, upon which is evidence of former sea levels. There is a boulder-beach deposit at 37–46 m, and fossiliferous beach rock deposits below 20 m. Most of the evidence supports a last interglacial sea level at 22 m above present sea level, indicative of extensive, more recent uplift of Tasmania. Fossil dunes of Pleistocene age occur on top of cliffs of 30–52 m high, and Holocene dunes are well developed on the eastern coastline. The west coast of Tasmania returns to the exposed ancient rocks with beach embayments described for the southwestern shore. In several locations there are terrace features at about 20 m, thought to be indicative of the last interglacial sea level as identified above from King Island. Granite outcrops occur between sections of narrower Pleistocene and Holocene dunes. South of the largest granite area near Granville Harbour is a large embayment between the quartzite headland protecting Macquarie Harbour to the south and the first of several granite headlands to the north. The result is extensive sand accumulation, and a common location for whale

strandings. Ocean Beach is a very high energy beach, one of the strongest in the world (Short 2006), facing the full force of westerly winds over the long fetch across the Southern Ocean. Average waves are 3–5 m, occasionally reaching as high as 18 m. Behind Ocean Beach is an extensive Pleistocene and Holocene dune field, where dunes reach 145 m height, and active and Pleistocene dunes extend up to 6 km inland over a 30-km coastline. Pleistocene dunes indicate the same west-northwest to east-southeast trending orientation as described for northeastern Tasmania. There is evidence of extensive recent erosion of the Ocean Beach coastline (>Fig. 21.7.8), cutting into dunes with old soil horizons, which is thought to be caused by recent relative sea level rise. Macquarie Harbour is a large (276 km2), almost enclosed harbour situated in a structural depression at the central west coast of Tasmania. The depression extends inland to the south, composed of Tertiary sediment of quartzite gravels and sands, and raised terraces indicating subsequent uplift of SW Tasmania. Two contrasting rivers discharge into Macquarie Harbour: the Gordon River from World Heritage Areas to the south, and the King River from the highly-disturbed Mount Lyell mining area to the west. Between 1916 and 1994, 97 million tons of tailings were discharged into the Queen River, a tributary of the King, and this resulted in building of a 2.5 km2 delta into Macquarie Harbour through the twentieth century. Deposits extend to a depth of 20 cm over most of the northern half of Macquarie Harbour (Locher 1997). ⊡⊡ Fig. 21.7.8 Erosion of dune coastline at Ocean Beach, Western Tasmania. (Courtesy M.-B. di Folco.)

Tasmania

Further south on the west coast of Tasmania resistant Pre-Cambrian quartzites form a broadly linear cliffed coastline controlled by a strike fault. This is indented by several drowned river valleys to form the Port DaveyBathurst Harbour estuarine basin of 130 km2. The coastline has cliffs and headlands 50–100 m high, some with cliff-top dunes, and seems to have been stable in Holocene times, despite high energy conditions.

References Bowden AR (1983) Relict terrestrial dunes: Legacies of a former climate in coastal northeastern Tasmania. Z Geomorphol Suppl. (Bd) 45:153–174 Bowden AR, Colhoun EA (1984) Quaternary emergent shorelines of Tasmania. In: Thom BG (ed) Coastal geomorphology in Australia. Academic Press, Sydney, pp 313–342

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Brothers N, Pemberton M, Pryor H, Halley V (2001) Tasmania’s offshore islands: seabirds and other natural features. Tasmanian Museum and Art Gallery, Hobart, 643p Davies JL (1985) Tasmania. In: Bird EC, Schwartz M (eds) The World’s Coastline. Van Rostrand Reinhold, New York, pp 975–979 Davies JL, Hudson JP (1987) Sources of shore sediment on the North Coast of Tasmania. Pap Proc R Soc Tasman 121:137–150 Edgar GJ, Barrett NS, Graddon DJ (2000) A classification of Tasmanian estuaries and assessment of their conservation significance: An analysis using ecological and physical attributes, population and landuse. Report to environment Australia from Tasmanian parks and wildlife service Kershaw RC, Sutherland FL (1972) Quaternary geomorphology of Flinders Island. Rec Queen Vic Mus Launceston 43:1–28 Locher H (1997) Mount Lyall remediaton. Sediment transport in the King River, Tasmania. Supervising Scientist Report, Canberra 120:124 Short AD (2006) Beaches of the Tasmanian coast and islands: a guide to their nature, characteristics, surf and safety. Coastal studies Unit and Surf Life Saving Australia Ltd, University of Sydney, 353p

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22.1 New Zealand

Terry Healy

1.  Introduction The diversity of coastal geomorphology and sedimentary deposits on the 18,000 km New Zealand coastline has evolved from a combination of geological structure, tectonic and seismic history, lithology, a mid-latitude oceanic setting for wave and tidal processes, Pleistocene events, and climatic influences. 1. Geological structure and lithology. A series of lineal axial ranges comprise the essential structural backbone of New Zealand. In the North Island these consist of lightly metamorphosed and intensely jointed Mesozoic greywacke, and in the South Island, of greywackes and metamorphosed Palaeozoic schists of Otago and the alpine fault zone through to gneiss and granodiorites of south Westland. On-lapping the axial ranges are ­younger Tertiary rocks, predominantly soft clay-rich siltstones and some limestones, while ancient volcanic mounds punctuate the landscape. In the central North Island the active Taupo Volcanic Zone stretches from Lake Taupo to beyond the Bay of Plenty coast. The northern North Island exhibits evidence of back arc andesitic and acid volcanism, with basaltic volcanism around Auckland and Northland. This diversity of lithology provides a wide variety of topographic forms and sediment for ­distinctive beach types and coastal landscapes. 2. Influence of structure and Pleistocene events. New ­Zealand sits astride a plate margin where the coastal geomorphology closely reflects the adjacent lithology as well as seismic events. Mesozoic and Tertiary orogenies created the basic axial ranges, the backbone of the country. Ongoing plate margin evolution saw the development of the transcurrent Alpine Fault along which active earthquakes and tectonic dislocations have influenced the coastal geomorphology by creation of drowned transverse fault aligned valleys (rias) of the Marlborough Sounds and uplifted coastal terraces around Wellington (Cotton 1942, 1974). Subduction of the Pacific plate under the central North Island results in a number of geologic processes ­including

the active Taupo Volcanic Zone, rapid tectonic uplift, ­folding, faulting and slumping with associated high seismic activity. These seismic and tectonic (uplift and subsidence) processes in the geological evolution, which are continuing at  present, have occurred on both a regional and local scale (Berryman and Hull 2003). They enhance denudation of the hilly and highland catchments, providing high sediment loads to the coasts. Regional subsidence has helped create the rias of the Marlborough Sounds, while regional uplift in the central North Island has resulted in incised rivers without significant estuaries (e.g. Waitara River) cut into coastal terraces. Events in the Pleistocene also played a major role in moulding the modern coastal landscapes. Severe alpine valley glaciation in the south Westland gneissic province produced glaciated U-shaped valleys, which upon drowning in the Holocene transgression have become classical fiords. The intense physical weathering of the fractured greywackes and schists, associated with the glaciations in the Southern Alps provided outwash sands and gravels to build up alluvial fan deposits, the most striking example of which is the Canter­­bury Plains. These subsequently became blanketed in loess. Periglacial weathering action in the ranges of the North Island assisted in delivering a large volume of gravel sediment to the lowland plains and coast, notably in Hawkes Bay. Of course at that time the lowland plains extended 100–120 m below present sea level, so that many sand deposits were reworked over the inner continental shelf as sea level rose with the post-glacial transgression. 3. Influence of volcanic acitivity. Both modern and ancient volcanic influences are evident in the coastal geomorphic landscape. Eroded caldera in the ancient volcanic mound of Banks Peninsula today contain the harbours of Lyttelton and Akaroa. Volcanic deposits have also had an important influence in the central North Island west coast where erosion of lahar flows provides both cobbles and boulders for veneer beaches over eroding shore platforms, as well as the distinctive black titano-magnetite heavy mineral sand for the extensive sand littoral drift systems of the North Island west coast. Propylitised Miocene andesites and

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_22.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Pliocene rhyolites and ignimbrites comprising the ­Coromandel Ranges result from back arc volcanics and have provided distinctive mineralogy for the beach sands of the embayed east Coromandel Peninsula. Around the Auckland isthmus some of the numerous Pleistocene basalt volcanic cones are now eroded headlands, with the outstanding landscape feature being the recently formed island of Ran­gitoto which erupted as Hawaiian-type fluid basalt flows, and is surmounted by a scoria cone crater active as recently as 600 years ago. Pleistocene and modern active volcanism from the Taupo Volcanic Zone has demonstrably influenced coastal evolution in the Bay of Plenty, where denudation of the acid volcanic tephras has provided large volumes of sand to the littoral zone. This has accumulated in the broad embayments to create extensive Holocene dune ridge progradations. The eruption of Mt Tarawera (1886), which blanketed the surrounding landscape with pumicious deposits, provides detailed evidence of the impact of a  large influx of sandy sediment into the coastal littoral system, and the airfall ash deposit on the coastal dunes allows ­reliable dating of progradation rates (Pullar and Selby 1971). The ability to use the technique of ­tephrachronology has allowed reliable dating of Pleistocene and Holocene dune and barrier deposits and interpretation of rates of geomorphic evolution along both the west and east coasts of the North Island. 4. Wave and climate influences. The mid-latitude oceanic islands of New Zealand are dominated by southwesterly waves originating from the Southern Oceans, and driven by the prevailing southwesterly winds. These waves, with a typical period of about 8–11 s and height of 1–3 m, drive littoral drift systems within regional and local coastal compartments along the west coasts. On the east coast of the South Island and the SE coast of the North Island the wave climate is likewise predominantly swell from the southern oceans. North of East Cape (North Island), the embayed coast (a ‘lee coast’ from the prevailing westerly winds) is sheltered from the southerly swells and is subject to a mixture of distant Pacific swells and local storm and wind generated waves, typically with periods of 7–9 s and heights of 0.5–1.5 m from variable onshore directions. Episodic storms, especially decaying tropical cyclones from the north, generate an important component of the wave climate that cause dune erosion problems (Hume et al. 1992). The El Nino-Southern Oscillation plays a major role in  the beach geomorphology and coastal oceanography.

In El Nino conditions New Zealand experiences continual 15–20 knot winds from the southwest. This causes a surfeit of orographically induced rainfall and flooding, particularly in Westland of the South Island, and enhancing sediment supply to the coast, while the constant onshore wind stress induces downwelling on the coast and inner shelf. But on the east coasts, particularly of the North Island, the predominant offshore wind stress induces drought conditions, onshore transport of nearshore sands, and inner shelf oceanographic up-welling. As a result the east coast beaches become generally well accreted in El Nino conditions and the coastal upwelling brings nutrient rich ocean waters to the inner shelf, which enhances conditions for algal and phytoplankton blooms in the nearshore coastal waters and harbours – a condition which severely impacts on fisheries and aquaculture industries. During the La Nina phase the opposite conditions occur, with more frequent, stronger northeasterly winds causing erosional tendencies for the northeast coast beaches, but upwelling induced algal blooms on the west coast. 5. Tidal effects. Tides are of the semi-diurnal type. Walters et al. (2001) calculated the tidal amplitudes around the oceanic islands of New Zealand based upon the predominant M2 constituent, to produce an approximate mean tide range. Based upon the tide gauges at the standard ports the average tide range is 1.8 m, with spring means of 2.1 m and neap means of 1.4 m. Microtidal ranges (3.5 m mean tide) occur in Tasman and Golden Bays in the northern South Island, and along the central west coast (e.g. Onehunga, Port Taranaki) and exhibit a mesotidal (>2 to Fig. 22.1.1), which in some Pleistocene dune deposits is of sufficient concentration to mine (Waverly, Taharoa), and in recent years has been prospected out on the continental shelf. There is a general reduction of heavy mineral concentrations away from Taranaki, both north and south. South of Cape Egmont, wave refraction induces a southeastward moving drift, but north of Cape Egmont there is a consistent northward drift to North Cape. Beaches along this sector typically exhibit a dissipative morphodynamic regime. Some of the littoral system is backed by cliffed

⊡⊡ Fig. 22.1.1 Boulder and cobble beach surmounting a broad shore platform cut in the lahar deposits that flowed from the Mt Taranaki. Note the concentrations of Titanomagnetite black “ironsand”.

  22.1

coast but the high energy waves also move sand on the submarine cut platforms of the inner shelf. The coastal sector centred on Taranaki is also undergoing rapid tectonic uplift forming coastal terraces, so that rivers, such as the Waitara, are actively downcutting and no large estuaries occur. In north Taranaki the Tertiary siltstones outcrop at the coast giving spectacular rapidly receding cliffs, with a basal abrasion notch. North of Kawhia Harbour a number of large shallow barrier enclosed estuarine lagoon harbours occur (Aotea, Raglan, Manukau, and Kaipara). Each is a tidal inlet, but unusual in that the barriers enclosing the harbours comprise Pleistocene dunes up to 300 m high, and depending upon tidal discharge relative to wave exposure, these tidal inlets tend to have well formed fan shaped ebb-tidal deltas (Hume 2003). Because of the high tidal ranges the ­currents in the inlet gorges tend to be strong – as much as 4 m/s. 2. Golden and Tasman Bays sandy compartments.  Loca­­ted at the northwest of the South Island between the Marlborough Sounds and Farewell Spit are two large embayed sandy compartments. Three features are remarkable here, namely the existence of a low dune ridge, Holocene sandy barrier island (>Fig.  22.1.2), and the 13 km long Nelson boulder barrier spit (>Fig.  22.1.3), which encloses a shallow muddy lagoon of Nelson Haven. About half of the enclosed area has been reclaimed, and the region possesses one of

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New Zealand

⊡⊡ Fig. 22.1.2 Rabbit Island, one of only two Holocene dune ridge barrier islands in New Zealand, adjacent to the Waimea inlet (far side of the island), which has the highest tidal range in New Zealand (about 4.5 m).

⊡⊡ Fig. 22.1.3 The 13 km long Nelson boulder barrier spit is a globally unique coastal landform enclosing Nelson Haven l­ agoon.

the highest tidal ranges in the country. At the extreme northwest is the large 30 km long Farewell Spit formed from the northwestward moving littoral drift fed by the sand component of the rivers discharging along the South Islands west coast. 3. Mixed sand-gravel sector of the South Island West Coast. Short steep gradient braided rivers draining from the mountainous Southern Alps provide abundant gravel and sands to the coast. Waves from the southern ­oceans drive a northward littoral drift.

­ eaches are usually mixed sand-gravel type, but some B cobble–­boulder ­beaches occur, and the wave climate is medium-to-high energy. As in the North Island west coast littoral system, sand evidently bypasses rocky bluffs, and ultimately feeds the large sand deposits of the 30 km Farewell Spit at the northern tip of the South ­Island. 4. The fiord coast of South Westland. The southwest tip of  New Zealand comprises a gneiss and granodiorite massif standing up to 3,000 m above sea level. At the

New Zealand

coast precipitous slopes drop down to, and below, sea level. Some boulder beaches occur but the coastline is predominantly cliffed. Pleistocene valley ­glaciations have carved a spectacular landscape of classical Ushaped valleys, forming classical fiords where they have become flooded at the coast (e.g. Milford Sound, Doubtful Sound). The location in the zone of ‘roaring forties’ and high topography means that the region receives very high precipitation, which provides fluxes of fresh water to drive a fiord-type estuarine salinity structure and circulation. As the area comprises national park the natural environment of this spectacular forested coastline remains in an essentially pristine condition. 5. The Southland-Otago cliffed and mildly embayed coast. The Southland coast varies between rocky sectors, low cliffed terrace coast, and embayed lowlands, with some gravel beaches. The low hills of southeast Otago result in a mainly cliffed coast with pocket sand beaches where rivers reach the sea. The region receives large swells from the southern oceans, and the beaches typically exhibit high energy morphodynamic conditions. Otago Harbour is a ria which has become largely infilled. 6. The Canterbury mixed sand-gravel coast. This is a distinctive and unusual littoral drift coastal sector formed from high energy storm and swell wave processes eroding the alluvial outwash gravel deposits comprising the Canterbury plains (Kirk 1980; ⊡⊡ Fig. 22.1.4 Storm-emplaced gravel beach attached to the southern side of Banks Peninsula.

  22.1

­Shulmeister and Rouse 2003). The result is a well formed steep berm beach comprised of mixed sand gravel sediment which are fed partly by direct erosion of the outwash gravel deposits and partly from episodic inputs from the several large braided gravel river systems of the Canterbury plains. Along sectors of the coast the active synoptic beach berm is surmounted by higher storm berms which reach almost 20 m above sea level adjacent to Banks Peninsula (>Fig. 22.1.4). Erosion of the outwash gravel deposits has resulted in low coastal cliffs behind the beach along the southern sector. South of Banks Peninsula backing dunes are rare, but they do occur on the sector north of the Peninsula. Banks Peninsula itself comprises a Miocene andesitic volcanic mound with eroded calderas creating present day harbours. Eroded loess deposits blanketing the hills have provided muddy infill sediment within the caldera harbours (>Fig. 22.1.5). 7. The North Canterbury and Marlborough Sounds ria coast. Along the north Canterbury coast the Kaikoura Ranges meet the sea resulting in a narrow coastal strip. At the northern tip of the South Island, fault aligned valleys, transverse to the coastal alignment, result in classical rias known as the Marlborough Sounds. The structural alignment and transcurrent faulting have created long linear rias (>Fig. 22.1.6). The steep hillslopes dip below the waterline with only ­minimal interruption to the profile and there is relatively little ­infilling

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New Zealand

⊡⊡ Fig. 22.1.5 Lyttelton Harbour is a drowned caldera in the volcanic mound of Banks Peninsula. Erosion scars expose the loess blanketing the volcanic rocks. This loess produces muddy sediment within the harbour.

⊡⊡ Fig. 22.1.6 Port Underwood, a fault-aligned ria coast on the Marlborough Sounds.  

s­ edimentation except at the valley heads, thus the rias remain deep. Wellington Harbour on the  southern North Island is likewise part of the ria coast. 8. The Eastern plate margin coast. Extending from the Wairarapa and Wellington to the East Cape is a rocky, eroding, high energy, mainly cliffed coast. This coast is seismically active and undergoing rapid uplift due to subduction processes at the plate margin. The general NE-SW alignment is broken by two large embayments,

Poverty Bay and Hawke Bay. Hawke Bay is subjected to high muddy river discharges to produce muddy sediment at shallow depths, but the rivers here also deliver greywacke from the central axial ranges, so that the beaches are predominantly composed of mixed sand and angular greywacke gravel. Resulting from the Napier earthquake of 1931, a large area of shallow sea floor and harbour was uplifted about 2 m to create new coastal lowland backed by degraded cliffs where marine processes are no longer active (>Fig. 22.1.7).

New Zealand

Elsewhere the coastline is surmounted by veneer sandy pocket beaches in shallow embayments, and often backed by raised marine cut terraces. The Poverty Bay lowlands comprise intercollated dune ridges and alluvial deposits. The high discharges from the Waipaoa River eject considerable muddy suspended sediment from the soft Tertiary mudstone catchments into the nearshore and continental shelf. Thus paradoxically, this coast is high  energy but possesses very muddy sediment at shallow depths due to ⊡⊡ Fig. 22.1.7 The sea floor of the Ahuriri estuary emerged during the 1931 Napier Earthquake.

⊡⊡ Fig. 22.1.8 The white siltstone cliffs of Young Nick’s Head and the mouth of the Waipaoa River, which provides sand for the progradation of the coastline, and mud to the nearshore and shelf.

  22.1

the high mud supply. Extensive shore ­platforms and buttressed cliffs are cut into the Tertiary siltstone ‘papa’ rocks. These present in a spectacular manner around Gisborne (>Fig. 22.1.8). 9. The Northeast embayed lee coast. Between East Cape and North Cape the coast is characteristically ­indented, including several large embayments with many islands (e.g. Bay of Islands, Hauraki Gulf). This is a coast

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New Zealand

which has evolved with predominant offshore winds in a swell wave environment. The Bay of Plenty contains the longest littoral system of the northeast region, and is dominated by low refracted swell waves of period 7–9 s and height 0.5–1.0 m. The bay is notable for containing the largest Holocene progradational extent of dune sands, 9 km in the Rangitaiki Plains, and the youngest barrier spit, Ohope Spit, which is less than 2000 years old. There are numerous shallow estuarine lagoons such as Maketu, and also the largest barrier island (Matakana Island) in the country (>Fig.  22.1.9). The extensive Tauranga Harbour, geomorphically a barrier enclosed estuarine lagoon, is about 70% exposed intertidal flats at low tide (Healy et al. 1996), and possesses a unique enclosing barrier system comprising Holocene barrier tombolos and an intervening barrier island. The surfeit of sand in the Bay of Plenty system is derived from the adjacent Taupo Volcanic Zone, having been brought to the coast during the Pleistocene sea level low stand, and reworked across the shelf during the Holocene post-glacial transgression, with continuous additions from

some of the larger rivers (Healy and Kirk 1982). Elsewhere in the less sand rich hilly coasts, Holocene sandy dune ridge barrier spits have built across the mouths of valleys and bays (>Fig. 22.1.10). The Firth of Thames graben is notable for containing the muddiest coast in New Zealand, with deep Holocene mud deposits comprising broad intertidal flats up to 4 km wide, and a well developed shelly chenier ridge plain (>Fig. 22.1.11). Separating the large sandy embayments the coast is largely steep and cliffed, with small bays and pocket beaches, as for example along the Coromandel coast, around the Waitemata harbour, and much of the Northland coast. The embayments comprise compartmentalised sandy littoral drift systems, some of which may be quite extensive (e.g. Bay of Plenty, Bream Bay, Doubtless Bay, Pakiri-Mangawhai), and contain extensive Holocene dune ridge sandy barrier progradation systems, the sand originally being derived from the Taupo Volcanic Zone as ancient Waikato River deposits (Schofield 1970). The typically hilly catchments and active erosion of the soft Tertiary lithologies means that the drowned river ⊡⊡ Fig. 22.1.9 The Mount Maunganui tombolo and Matakana barrier island, formed of Holocene sandy deposits and enclosing the estuarine lagoon of Tauranga Harbour. (Courtesy Port of Tauranga.)

New Zealand

  22.1

⊡⊡ Fig. 22.1.10 The Holocene barrier spit of Pauanui constructed across the Tairua valley on the east Coromandel Peninsula.

⊡⊡ Fig. 22.1.11 The broad mudflat and shelly chenier ridge plain at Miranda, Firth of Thames, with a welldeveloped estuarine turbid fringe.

valleys have to a large extent become infilled with sediment, resulting in shallow lagoonal harbours with extensive intertidal flats. At the mouth of many former drowned river valleys are Holocene sandy dune ridge barrier spits (e.g. Pauanui, Matarangi), which enclose shallow infilling estuarine lagoons. Within the Hauraki Gulf the soft Miocence alternating sandstone-siltstone

flysch cliffs have undergone rapid erosion in the Holocene and today display active cliffs and extensive shore platforms up to 400 m wide (>Fig. 22.1.12) (Moon and Healy 1994). In the far north Parengarenga Harbour is also an estuarine lagoon, dominated by white quartz sands that probably originated from erosion of strongly podzolised soils.

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New Zealand

⊡⊡ Fig. 22.1.12 The cliff and wide intertidal shore platform cut into dipping Miocene flysch deposits, Whangaparoa Peninsula.

References Berryman K, Hull A (2003) Tectonic controls on late quaternary shorelines: A review and prospects for future research. In: Goff J, Nichol S, Rouse H (eds) The New Zealand Coast. Dunmore Press, Palmerston North, pp 25–58 Cotton CA (1942) Geomorphology. An introduction to the study of landforms. Whitcombe and Tombs, Christchurch, 505 pp Cotton CA (1974) Bold coasts. Reed, Wellington, New Zealand 354 pp Healy TR, Cole R, De Lange W (1996) Geomorphology and ecology of New Zealand shallow estuaries and shorelines. In: Nordstrom KF, Roman CT (eds) Estuarine shores: evolution, environments and human alterations. Wiley, New York, pp 115–154 Healy T, Kirk RM (1982) Coasts. In: Soons JM, Selby MJ (eds) Landforms of New Zealand. Longman Paul, Auckland, pp 81–102 Hume TM, Bell RG, De Lange WP, Healy TR, Hicks DM, Kirk RM (1992) Coastal oceanography and sedimentology in New Zealand, 1967–91. N Z J Mar Freshwater Res 26:1–36

Hume TM (2003) Estuaries and tidal inlets. In: Goff J, Nichol S, Rouse H (eds) The New Zealand Coast. Dunmore Press, Palmerston North, pp 191–214 Kirk RM (1980) Mixed sand and gravel beaches: morphology, processes and sediment. Prog Phys Geog 4:189–210 Moon V, Healy T (1994) Mechanisms of coastal cliff retreat and hazard zone delineation in soft flysch deposits. J Coastal Res 10:663–680 Pullar AS, Selby MJ (1971) Coastal progradation of Rangitaiki Plains. N Z J Sci 14:419–434 Schofield JC (1970) Coastal sands of Northland and Auckland. N Z J Geol Geophys 13:767–824 Shulmeister J, Rouse H (2003) Gravel and mixed sand and gravel systems. In: Goff J, Nichol S, Rouse H (eds) The New Zealand Coast. Dunmore Press, Palmerston North, pp 143–162 Walters RA, Goring DG, Bell RG (2001) Ocean tides around New Zealand. N Z J Mar Freshwater Res 35:567–579

22.2 The Chatham Islands

Eric Bird

1.  Introduction The Chatham Islands (>Fig. 22.2.1) lie about 856 km east of Christchurch at latitude 44° S and longitude 176° 30' W. They are thus strictly in the western hemisphere, but the International Date Line has been moved east to avoid confusion. The three largest islands are Chatham Island (90,000 ha), Pitt Island (6,190 ha), and South East Island (218 ha), and there are six smaller islands and many tall cliffy islets and steep rocky stacks. About 40 km to the east are the high stacks known as Motuhara (Forty Fours) Rocks (43° 58' S, 175° 30' E), then about 8,000 km of open ocean to the southern coast of Chile. The islands were sighted by William Broughton in 1791, who named them after his naval ship, Chatham. They had already been occupied by a Polynesian (Moriori, ­pre-dating the Maori) population for several centuries. They became part of New Zealand in 1901. The islands stand above the Chatham Rise, an elongated shelf platform (submerged land area) 90–150 km wide at depths of up to 600 m, extending eastward from the Banks Peninsula (Campbell et al. 1993, Campbell 1996). They are the emergent part of the Chatham Rise, elevated by tectonic uplift and volcanic activity. The northern part of Chatham Island is generally low-lying and gently undulating, and on the shores of the NW and NE peninsulas there are hard, grey, quartz-veined foliated rocks (Chatham Schist) similar to those seen in Otago, formed by the metamorphism of Permo-Triassic sedimentary rocks. It is unusual to find continental rock formations of this kind far out on the Pacific Ocean islands, and they also form the Forty Fours to the east. There are Cretaceous to Cainozoic marine limestones and sandstones and Quaternary dunes and dune calcarenite in the centre and north of Chatham Island. There are also small conical volcanic hills, 100–150 m high, of Upper Eocene and Lower Oligocene basalt, in a line that runs east from Mount Differbach (Hewokama) to Hokopoi, Motupor­ oporo, and Korako. There are sediments of volcanic origin, including tuffs (consolidated ash) and agglomerates that were deposited in the sea, and contain marine fossils. The embayed coast has beaches of white or yellow calcareous

sand, backed by grassy dunes. The central part of Chatham Island is an isthmus of dunes and dune calcarenites over Tertiary limestones, with Te Whanga Lagoon (16,190 ha) on its eastern side. This brackish lagoon has an intermittent entrance through a dune-capped barrier on the east coast north of Owenga and limestone on its western shores. There are several other shallow lakes and lagoons. The southern part of Chatham Island is higher, on Upper Cretaceous volcanic rocks, rising to a steep ­southern coast with cliffs up to 250 m high, exposing northwarddipping basaltic lava, tuffs, and agglomerates. Similar rocks are found on Pitt Island, a dissected remnant of a broad volcano centred on Pitt Strait. Pitt Island and the smaller islands and stacks are bordered by precipitous slopes and cliffs. There are numerous trachyte dykes, and many small folds and faults in the volcanic rocks. Peat deposits are extensive, up to 10 m thick on the southern uplands of Chatham Island and on Pitt Island, in areas typically dominated by rushes and tarahinu (Dracophyllum arboretum) scrub. They began to accumulate between 30,000 and 40,000 years ago, and contain thin layers of volcanic ash from the Lake Taupo eruptions, the thickest (16 cm) from the Oruanui event about 23,000 years ago. There are Quaternary benches cut by marine processes at various higher sea levels up to 200 m, notably on South East Island. They are typically veneered with sand, capped by peat. The climate is cool, cloudy, and windy. At Waitangi mean monthly temperatures rise from 7.5°C in July to 14.3°C in January, and average annual rainfall is 500–1,000 mm, rising to over 1,250 mm in the higher country to the south. The prevailing winds are southwesterlies. The mean spring tide range is about 0.9 m at Waitangi. The southern and western coasts receive strong southwesterly ocean swell, while generally gentler Pacific swell arrives from the east. Occasionally, large tsunami waves reach the Chatham Islands: they were recorded here in 1868, 1934, and 1964. Earthquakes, related to tectonic movements at the eastern end of the Chatham Rise, occur rarely. Exotic rock types, found particularly on the shores, include obsidian and greenstone implements from New Zealand and some erratics, probably borne by icebergs from Antarctica;

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_22.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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The Chatham Islands

⊡⊡ Fig. 22.2.1 The Chatham Islands. CY – Cape Young, CP – Cape Pattison, PH – Port Hutt, PB – Petre Bay, W – Waitangi, PD – Point Durham, CL – Cape L’Eveque, S –The Slump, MP – Manukau Point, HB – Hanson Bay, TWL – Te Whanga Lagoon, PM – Point Munning, FP – Flower Pot Bay, MI – Mangere Island, SEI – South East Island. (Courtesy Geostudies.)

i­ cebergs were seen occasionally off the Chatham Islands until around 1950. In Waitangi Bay lumps of Welsh slate and English flint derived from ships’ ballast are sometimes washed ashore, often adhering to kelp holdfasts. Surpris­ ingly, in view of the islands’ remoteness, maritime litter is strewn on many beaches and shores.

2.  Chatham Island The coastal features of Chatham Island are described and illustrated in sequence, beginning at Cape Young in the northwest and proceeding down the west coast, along the south coast, up the east coast, and along the north coast. Cape Young is bordered by cliffs and rocky reefs cut in Pliocene volcanic deposits. On Momoe-a-Toa Point to the south the cliffs expose the Lower Pliocene Momoe-aToa Tuff Formation with shell beds including fossil scallops. The cliffs are dissected by caves and blowholes. At Tioriori the mouth of Tutuiri Creek is bordered by cliffs cut in late Cretaceous to Palaeocene Greensands, the Tioriori Group, marine glaucontic, and phosphoritic

sediment with fossiliferous horizons. Northwesterly swell breaks on Maunganui Beach, which is backed by three parallel dune ridges showing a zonation from grasses landward to scrub and woodland. Mount Maunganui (178 m) is a Pliocene volcanic plug. The coast curves out to a low headland at Cape Pattison, fringed by rocky outcrops of grey foliated quartzveined Chatham Schist. On the western side is Waitangi West Beach, a curved beach shaped by westerly swell, and the schist reappears on the shore to the south and extends round the Point Hutt Peninsula to Petre Bay. South from Point Alison the rocky coast is indented by narrow inlets at the mouths of parallel valleys drained by streams such as Tangoio Creek and Rongotea Creek. Te Raki Point and Point Somes are broader headlands on either side of Te Koparuparu Bay, and at Point Somes the rocky coast swings eastward and becomes embayed. Waipurua Bay and Ocean Bay are followed by Port Hutt (Whangaroa Harbour) (>Fig. 22.2.2), and a series of south-facing bays, which are submerged valley mouths fringed by schist shores. At Ohira the schist is interrupted by columnar Upper Cretaceous basalt, dated at 79 million years, and there are bay-head dunes. At Paritu the rocky coast comes to an end and Long Beach, a sandy surf beach, backed by dunes bearing ­marram grass, extends behind Petre Bay (>Fig. 22.2.3). Parabolic dunes and blowouts have been shaped by the  prevailing southwest winds. There are several dune lakes inland, such as Te Roto (Tennants Lake) and Lake Marakapia, where the Holocene dunes extend over Pleistocene dune calcarenite and a basement of Tertiary limestone. At Big Bush a quarry shows red tuffaceous clay  over Te Whanga Limestone, the contact showing ­evidence of baking beneath volcanic ash. At Red Bluff the  beach is interrupted by a sector of cliffs up to 70 m high, cut in the Red Bluff Tuff Formation, consolidated volcanic ash and scoria, containing bryozoan and molluscan fossils. To the south the sandy beach resumes, backed by lower dunes and curving round into Waiutangi Bay. Waitangi stands at the mouth of the Nairn River. There are Upper Pleistocene to Holocene shell beds in the bluff to the east of Waitangi, and the cliffs in weathered tuff rise past the wharf (>Fig. 22.2.4) and out to Hanson Point and Point Webb. There are several lava mounds on the Waitangi Peninsula. At Point Weeding the cliffs are cut in Red Bluff Tuff (>Fig. 22.2.5). The coast turns south, then southwest, and gentle slopes descend to cliffs cut in tuff. There are several trachyte dykes. Basalt emerges to form kelp-edged shore platforms and reefs (such as Heaphy Shoal) along the coast to Point Durham, where a lagoon is impounded by a rocky ridge.

The Chatham Islands

  22.2

⊡⊡ Fig. 22.2.2 Port Hutt, with a curved sandy beach at its head. (Courtesy Geostudies.)

⊡⊡ Fig. 22.2.3 Petre Bay, with ocean swell arriving on a grey sandy beach backed by calcareous dunes. (Courtesy Geostudies.)

At Point Durham the coast turns southward, and basalt outcrops in cliffs and a rocky shore. The coastal slopes rise and steepen, and are incised by the valleys of incised streams, such as Little Awototara Creek. Rocky reefs of dark basaltic lava are prominent on Point Gap and Otawae Point, and there are columns of diorite at Te Whakahewa (The Horns). At Cape L’Eveque the south coast of Chatham Island begins, with cliffs below a steep escarpment on northward-dipping basalt rising to 200–250 m. The uplands carry a peat mantle. Valleys cut by incised streams end in

coastal waterfalls east of Green Point, and at Cascade Gorge waterfalls have receded inland behind a deepened valley. Houruakopara Island, 37 m high, is a residual left by cliff recession, and Ka Oreao Point projects as a bold interfluvial ridge, rising to 119 m. To the east is The Slump, a coastal landslide backed by steep bluffs with streams cascading down the slope (>Fig. 22.2.6), and cliffed at the base behind a gravel beach; it is fairly stable and largely vegetated. The steep escarpment coast with high cliffs and waterfalls continues NE past Rangaika (264 m), with more

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The Chatham Islands

⊡⊡ Fig. 22.2.4 Cliffs in weathered tuff at Hanson Point, looking towards the Waitangi Wharf. (Courtesy Geostudies.)

⊡⊡ Fig. 22.2.5 Cliffs cut in weathered tuff overlying solid tuff, in which a shore platform has been cut at Point Weeding. (Courtesy Geostudies.)

s­ ubsided sectors between upper and lower cliffs, then the basalt cliffs gradually diminish past Boundary (152 m) to Cape Fournier. As the coast turns northward the steep slope declines to only minor basal cliffs, and diminshes diminishes to the low headland of Manukau Point. Gentle slopes then descend to low cliffs on the north-facing coast past Cloughs Creek to Owenga village and Hawaiki Creek, where the black basalt outcrops come to an end. There are emerged Holocene shell beds, Moriori kitchen middens, and dumped scallop shells.

The low-lying east-facing coast of Hanson Bay has a long, sandy beach shaped by easterly swell and backed at first by cliffs cut into sandy humate (>Fig. 22.2.7 ). This is a black organic sand that evidently formed by peat accretion in accumulating sand dunes. The beach is dominated by black sand derived from the humate cliffs. The humate is capped by grey dune sand in which there are Moriori kitchen middens (>Fig. 22.2.8). To the north the humate cliff diverges inland behind parallel sand dunes, which continue to the entrance to Te Whanga Lagoon.

The Chatham Islands

  22.2

⊡⊡ Fig. 22.2.6 View westward across The Slump, showing the basal cliff. (Courtesy Geostudies.)

⊡⊡ Fig. 22.2.7 Cliffs cut in sandy humate north of Owenga. (Courtesy Geostudies.)

Te Whanga is a coastal lagoon formed where sandy barrier formations have been built up to link former islands in such a way as to enclose about 160 sq. km of brackish water (>Fig. 22.2.9). The lagoon is generally cut off from the sea by a shelly sand barrier, but after heavy rains and westerly gales an entrance opens temporarily. The entrance zone is a kilometre-wide gap in the grassy dunes, and consists of a wide bank of buff sand strewn with shells, rising 2–3 m above the mean level of the lagoon. Tokatapu Shoal, offshore to the SE, has probably

influenced the location of the lagoon entrance. The northern dunes are cliffed. On the inner side of the entrance to Te Whanga Lagoon rushy islands form a tidal delta, with shoals and a sandy arrow spit (Waikawa Island) pointing westward across the shallows towards Te Matarae Point. In 1988 an attempt was made to establish a permanent entrance when a unit of the New Zealand army corps excavated and blasted a channel, but it was soon sealed by sand deposition. To the north the inner shore of the dune-capped barrier is

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The Chatham Islands

⊡⊡ Fig. 22.2.8 Moriori kitchen midden above the humate cliffs north of Owenga. (Courtesy Geostudies.)

fringed by a series of elongated and arcuate sand bars, thought to be migrating eastward on to the shore, and a wide marshy fringe. On the northwestern shore of Te Whanga Lagoon there are outcrops of schist which give place to limestone cliffs and ledges towards Kainga Paheka (>Fig. 22.2.10). These may originally have formed when this coast was exposed to stronger wave action, before the formation of the enclosing eastern dune barrier, as waves on Te Whanga Lagoon are small. There are several sectors of chalky Te Whanga Limestone cliffs and bluffs, with some eroded stacks, separated by coves with beac hes of shelly sand. The Te Whanga Limestone is a Lower Tertiary (Palaeocene to Eocene) marine bryozoan and foraminiferal deposit containing echinoderms and black sharks’ teeth, found at Blind Jim’s Creek, north of Cattle Point. The broad low Karewa Peninsula runs out past Titirangi Point to Karewa Point, and round to Moutapu Point. At Titirangi Point and Moutapu Point the cliffs are cut in the Upper Pliocene Titirangi Sand Formation, a sandy sediment with shell fragments. Te Whanga Limestone outcrops resume in low cliffs southwest to Waikato Point and Motohou Point, with an outcrop of Eocene Te One bryozoan Limestone at Moeroa. Along the west coast of Te Whanga Lagoon there is evidence of Holocene emergence in the form of degrad­ ­ed cliffs, wide shelly beach ridges, and low sandy or rocky terraces. Offshore, karstic ribs of rock mark former limestone islands, as at Motuhinahina and Waipoua. In the absence of strong wave action, solution processes are effective, producing notches on outlying stacks (>Fig. 22.2.11). There are intervening sandy coves and some longshore

spits (>Fig. 22.2.12). South of the low rocky point of Te Matarae, with its irregular calcarenite stack, the bluffs expose Palaeocene Red Bluff Tuff, and at the southern end is a rushy delta at the mouth of the Awainmanga River. The lagoon water is much diluted by peaty stream inflow; the floor is sandy. In dry seasons the lagoon level falls and wide bordering sandy and muddy flats are exposed, but after heavy rains it rises until eventually there is outflow at the entrance. After episodes of strong northerly or southerly wind the lagoon waters oscillate, with seiches of up to 60 cm. Much of Te Whanga lagoon is shallow, but it is up to 10 m deep off Waikato Point, from where two deep channels run S and SSW. It is partially segmented by shallow zones at Karewa and Te Matarae. Shore vegetation along the east shore, and beyond Kahupiri Point, is dominated by rushes and some shrubs, but there are no mangroves. North of the lagoon entrance at Te Awapatiki the sandy shore of Hanson Bay is backed by parallel grassy dune ridges up to 20 m high and a series of small freshwater peaty lakes, including Long Pond, Lake Taia, Lake Kairae, Lake Makuku, and Lake Kaingarahu. The high foredunes continue past Hapupu, backed by Lake Rangitai, which is shallow with a broad, sandy fringe. The low, gently undulating hinterland has dune topography on Pleistocene aeolian calcarenite. The curving beach comes to an end as a rocky shore consisting of Chatham Schist, emerges east of Lake Rotorua, and low cliffs run out to Okawa Point. The bay to the north is backed by a beach with schist outcrops, then grassy dunes, and the northeastern promontory has rocky shores (Chatham Schist) out to Te Wakaru Island,

The Chatham Islands

  22.2

⊡⊡ Fig. 22.2.9 Te Whanga Lagoon. (Courtesy Geostudies.)

separated by a narrow channel. Point Munning, the NE point of Chatham Island, is also low and schist-fringed. A series of small coves with rocky shores runs west from Point Munning to the wider bay at Kaingaroa Harbour. This is backed by a sandy beach and an eroding dune cliff, with shore platforms cut across jointed Chatham Schist on either side. To the west are sandy bays backed by grassy dunes between low promontories of schist. Lake Pateriki is separated from the sea only by a bare sand bank, over which the sea occasionally washes. Chatham Schist

fringes promontories to Matarakau Point and forms French Reef, in the lee of which is the lobate Ocean Mail Point on a sandy shore backed by grassy dunes. This becomes a narrow dune-capped barrier at Pakauwera, with Te Whanga Lagoon on the southern side. There are peaty soils in dune swales and pebble beds beneath the dunes at Mananea. The sandy beach then curves out to Taupeka Point, where there is Cretaceous basalt on the shore, related to the Southern Volcanics. There is much seaborne litter on these northern shores. West of Taupeka

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The Chatham Islands

⊡⊡ Fig. 22.2.10 Limestone outcrops on the western shore of Te Whanga Lagoon. (Courtesy Geostudies.)

⊡⊡ Fig. 22.2.11 Mushroom stack near Waipoua, undercut by solution processes. The notch is at normal water level, Te Whanga Lagoon being at a low level when the photograph was taken. (Courtesy Geostudies.)

Point there are intermittent shore outcrops of Chatham Schist between sandy beaches backed by dunes disturbed by blowouts from the southwest. The shore platform cut in schist widens towards Cape Young.

3.  Pitt Island Pitt Island (Rangiauria) has mainly steep coastal slopes, with cliffs on the west and southeast coasts, rising to lush

grassland over peat and guano deposits. The western coast is higher, with rivers rising a short way inland and flowing down to the east. The island has rich soils formed on weathered basalt overlain by peat and guano. Motutapu Point at the northern end is a spur rising to 62 m. The marine Pliocene Whenuataru Tuff outcrops in cliffs here and to the south in Flower Pot Bay (Onoua), where it is interbedded with white shelly Pliocene Onoua Limestone. At Kahuitara Point in Flower Pot Bay the cliff is cut in Upper Cretaceous sandy tuff with ammonites and

The Chatham Islands

  22.2

⊡⊡ Fig. 22.2.12 Longshore spit formed by southward drifting of sand on the west coast of Te Whanga Lagoon, north of Te Matarae. (Courtesy Geostudies.)

belemnites. Tarawhenua Point is in the Red Bluff Tuff Formation with Pliocene Whenuataru Tuff, and there are basaltic lava cliffs on its southern side. Rabbit Island has no rabbits, but plenty of mutton bird burrows. On the west coast of Pitt Island Waihere Bay is backed by slumping slopes in the soft, sandy Middle Cretaceous Tupuangi Formation. There are numerous transverse faults, two of which delimit a bold cliff of Red Bluff Tuff with lenticles of Te Whanga Limestone protruding from the slumping slopes in the middle of the bay. Trachyte dykes form a bold promontory and Rangiauria Point has cliffs of Pliocene lava. The southernmost point, Murumuru (185 m), has steep cliffs with many caves and outlying high islets. On the southeast coast of Pitt Island high cliffs of basalt run behind Canister Cove (Waikokopu) and Glory Bay, and north of Waipaua a steep coast descends to lower cliffs. The volcanic rocks continue to Kahuitara Point, the eastern cape, but pass to Cretaceous sedimentary formations along the northeast coast to Tupuangi. Here streams descend into a lake that has an outlet to the coast. The steep basalt coast then returns, becoming cliffed to the north at Mototapu.

4.  Other Islands South East Island, rising to 224 m, has high cliffs in dipping coarse Pliocene volcanic sediment on its west coast and slopes descending steeply to an eastern bay below Rangatira. To the west of Pitt Island Mangere Island and its neighbours are the remains of a Pliocene volcano. The island has steep coasts with high cliffs, and a bouldery northeastern shore. A narrow strait separates it from Little Mangere Island (The Fort) to the west, which is bordered by lava cliffs rising more than 200 m above the sea. To the southwest is The Castle, a high narrow ridge rising to 156 m, flanked by cliffs in volcanic rock, and Sail Rock is a bold outlying stack to the west.

References Campbell H (1996) Geology. In the Chatham Islands: heritage and conservation. Canterbury University Press, Chri, New Zealand, pp 34–48 Campbell H et al (1993) Cretaceous-Cenozoic geology and biostratigraphy of the Chatham Islands, New Zealand, Vol 2. Institute of Geolog­ ical and Nuclear Sciences, Monograph, Wellington, New Zealand

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23.0 Pacific Ocean Islands – Editorial Introduction

Islands in the Pacific Ocean are of three kinds (Nunn 2005). Some are fragments of continental structures, such as New Caledonia, others are of volcanic origin, such as Easter Island, and many are coralline, such as the high limestone islands (emerged atolls) of the Loyalty Islands and the cays of Kiribati. The islands show the effects of submergence by the Holocene marine transgression, a brief phase of higher sea level in the mid-Holocene and subsequent emergence, complicated by neotectonic movements. Most of the islands lie in the SW Pacific, but the Galapagos, Clipperton, and Easter Island are volcanic islands rising from the East Pacific Ridge (>Fig. 23.0.1). > New Caledonia is an elongated mountainous island of mainly metamorphic rocks, including schists, peridotite, and serpentine, deeply weathered under a humid tropical climate. It has a steep, partly cliffed eastern coast, and a gentler western coast with several alluvial plains. There are segments of fringing coral reef and an encircling barrier reef. The uplands have been drastically modified by large-scale open-cast mining of nickel, which has generated massive quantities of rock debris, gravel, and sand that have spilled down mountain slopes into valleys, and been delivered to the coast by rivers. The outcome is pollution of the sea, damage to coral reefs, and augmented sedimentation on mudflats and mangrove swamps. Volcanic islands occur at the boundaries of tectonic plates, notably along the West Pacific zone that runs from New Zealand northward through the Kermadec Islands, the Solomon Islands. and New Guinea, through to Japan. They also follow migrating mid-ocean hot spots, as in > Hawaii. Some volcanoes are still active, as on the island of Hawaii, but most are dormant or extinct. The Society Islands show volcanoes in stages of dissection by runoff  and rivers. > Fiji has dissected volcanic formations

accompanied by marine sedimentary rocks, coral limestones, and alluvial deposits, notably on the Rewa, Singa­ toka, and Nadi deltas. In the Cook Islands emerged fringing reefs formed by a lowered sea level encircle volcanic cones. Coral reefs are widespread, especially in the West Pacific. They include atolls where a rock foundation, usually volcanic, has subsided, Bora Bora in the > Society Islands showing a stage when a central volcanic island has persisted, surrounded by a lagoon and enclosing coral reef. Uplifted coral forms high limestone islands, which occur where there has been disruption at converging plates: the Loyalty Islands, NE of New Caledonia, include raised atolls (Mare and Lifou) and a tilted atoll (Ouvea). The coral limestone forms cliffs, dissected by caves. There are often stairways of terraces indicating stages in uplift. An emerged coral platform surrounds the Isle of Pines, south of New Caledonia. Low islands have formed where sand and gravel eroded from coral reefs has been piled up above high tide level. Such cays are seen in Kiribati and the Marshall Islands. Some consist entirely of unconsolidated sediment, and are ephemeral or migratory. Beach rock (cay sandstone) and beach conglomerate are formed by cementation of coralline sand and gravel, and may be exposed by beach erosion. More durable cays, termed motu, have a core of emerged reef limestone. These low-lying islands are threatened by a rising sea level.

Reference Nunn PD (2005) Pacific Ocean islands, coastal geomorphology. In: Schwartz ML (ed) Encyclopedia of Coastal Science. Springer, Dordrecht, the Netherlands, pp 754–757

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_23.0, © Springer Science+Business Media B.V. 2010 (Dordrecht)

⊡⊡ Fig. 23.0.1 Pacific Ocean Islands. (Courtesy Geostudies.)

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23.0 Pacific Ocean Islands – Editorial Introduction

23.1  New Caledonia and the Loyalty Islands Jean-Michel Lebigre · Eric Bird  

1.  Introduction New Caledonia (area, 19,800 sq km) has a coastline about 2,254 km long. It is a mountainous island (rising to 1,629 m on Mont Panié) consisting mainly of crystalline schists with peridotite and serpentine rock. The east coast is generally steep, with only minor plains around river mouths, several of which open into deep rias; the west coast has more extensive plains, backed by a hilly hinterland, with numerous promontories and islands, and intervening embayments into which deltas have been built. Open-cast mining of nickel is extensive in the lateritised peridotite highlands of New Caledonia, and has added red clay, sand, gravel, and boulders to the natural sediment loads of many of the rivers, eventually accelerating the infilling of estuaries, the progradation of deltas, and the shallowing of nearshore waters. Evidence of such changes has been obtained by comparing nineteenth-century maps and charts with air photographs taken in 1954 and 1976 (Bird et al. 1984). Stevenson et al. (2001) have sought to elucidate Late Quaternary environmental change and human impacts in New Caledonia. The coastal climate has a dry winter and a hot wet summer when tropical cyclones are found. Nouméa has mean monthly temperatures of 19°C in July and 25°C in January, and an average annual rainfall of 1,083 mm. Southeasterly trade winds prevail. The coastal waters of New Caledonia are rich in coral reef formations. Encircling barrier reefs include the atolllike loops of D’Entrecasteaux and Great South reefs, and numerous knolls and patch reefs, some with sandy cays. Fringing reefs are best developed on west-facing shores (Chevalier 1973). Some have a shallow boat channel to the rear, and others are attached to the shore, the inner parts bearing algal deposits, sandy flats backed by derived white coralline beaches, or mangroves. Tide ranges are small, generally less than a metre. Nouméa has a mean spring tide range of 1.1 m and Houailou 1.0 m. Wave action is related to local, mainly southeasterly, winds, except where oceanic waves (southwest swell on the west coast, easterly swell on the east coast) enter through the passes that interrupt the ­barrier

reef, and strong wave action penetrates to adjacent shores. Cliffs are found only in a few sectors reached by strong wave action, as at Cape Ndoua and on the southeast of Ile Ouen in the far south of New Caledonia. Bold coasts are more extensive, usually with steep seaward slopes bearing a soil mantle and vegetation cover, and only limited basal cliffing. Mangrove swamps occupy more than 27,000 ha in New Caledonia. Narrow shore fringes are widespread, but the largest stands are in the Bay of Nehoué, and on either side of the Poum Isthmus in the northwest. In the sheltered situations where sediment is abundant, mangroves advance rapidly: up to 400 m between 1954 and 1976 on the shores of the Népoui delta, where red mud is accumulating as the result of mining operations upstream.

2. The Coastline of New Caledonia (La Grande Terre: the Main Island) Nouméa stands on the neck of an elongated branching peninsula that runs out to Ouen Toro. Small bays alternate with rocky shores. To the north is the wide bay, the Grande Rade, at the head of which is the Doniambo nickel smelter. The coastline hereabouts is largely artificial. A little to the south, Nou Island (125 m at Mount Tereka) has been attached to the mainland by an embankment. The peninsula of Ducos, to the north, and beyond it the Baie de Koutio-Koueta, have many headlands and small mangrove fringed bays. Offshore is the sheltering barrier reef. In the 20 km wide lagoon, there are many small cays on coral banks, several of which are marine protected areas (the Réserves Spéciales du Lagon Sud). In the Baie de Dumbéa, the Dumbéa River has built a large delta with zoned mangrove communities, typical of the west coast of New Caledonia, where the pioneer mangroves are colonizing an area of slow and regular terrigenous accretion (>Fig.  23.1.1). The Tontouta and Ouenghi rivers are building mangrove-fringed deltas into thisbay. The Baie de Gadji is swamp-fringed, there are ero­­ded  promontories (and islands, some linked by

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_23.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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New Caledonia and the Loyalty Islands

⊡⊡ Fig. 23.1.1 The advancing seaward edge of mangroves on the Dumbéa Delta.

⊡⊡ Fig. 23.1.2 On the Païta area, the sandy tombolo of T’Ndu connects Tiaré Peninsula with the small TNdu Island.

tombolos (>Fig.  23.1.2), and narrow sandy beaches behind fringing reefs. Farther north, the isleted and reef-protected Baie de St. Vincent has intricate outlines, with many headlands and bays, and steep slopes rising to high hinterland. The bay

receives terrigenous sands and clays from rivers and coralline sediment from the sea floor. Dune calcarenites formed during Pleistocene low sea level stages on the outer coasts of Puen, Hugen, and Mathieu islands, but these are otherwise rare in New Caledonia. The Tontouta River flows in through

New Caledonia and the Loyalty Islands

23.1

⊡⊡ Fig. 23.1.3 Ouenghi delta, Baie de St. Vincent. (Courtesy Geostudies.)

⊡⊡ Fig. 23.1.4 Cliff and shore platform on an exposed headland south of Onghoué.

mangrove-fringed distributaries in a delta in southeast, and the Ouenghi River arrives through a swampy delta to the north (>Fig.  23.1.3). Some mangrove areas have been adapted to shrimp (prawn) ponds. The clay plain behind the Bay of Déama carries sandy cheniers. The peninsulas and islands of the Baie de St. Vincent have steep coasts, cliffed on the seaward side with shore (abrasion) platforms (>Fig.  23.1.4), but the outlying barrier reef reduces wave

impact. Steep Ile Léprédour (>Fig. 23.1.5) and the Bouraké Peninsula are developed on Triassic Monotis Limestone, with the Baie de Saint Vincentto the north. The coast northwestward has a series of wide bays between promontories, and is largely reef-fringed. Near Bourail, the Néra River flows to a delta behind the Baie de Gouarou, with a shore consisting of a gently curving sandy beach shaped by southwesterly waves that come in through

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New Caledonia and the Loyalty Islands

a gap in the reefs. This is the seaward shore of a sandy barrier spit backed by mangroves. La Roche Percée has a prominent stack, La Bonhomme, cut in sandstone. Beach rock is exposed on the eastern shore of the adjacent Baie des Tortues, which has an emerged reef on the western side. To the northwest, extensive fringing reefs are backed by the calcareous sandy Plage de Poé with extensive beach rock and numerous parallel beach ridges. White coralline sand drifts on to the beach from the fringing reef, and locally beach rock is exposed. The limestone of Les Montagnes Blanches comes to the coast at Cap Goulvin. There is a contrast between the embayed Néra River delta and the more intricate delta of the Poya River, to the north, which has mangrove-­bordered lobate salients between distributary mouths opening into a bay, sheltered from strong wave action, with numerous shoals threaded by channels in the Passe des Rivières Prolongées. The bay is bordered by islands and peninsulas of Miocene limestone. Mangrove-fringed salt marshes and dry saline flats are extensive on these west coast deltas and plains, which are submerged by summer river flooding and cyclone surges, but desiccate in the long dry season. The marshes of Mara partly overlie a dead coral reef, and have saline flats backed by a clifflet a few centimetres high rising to savanna terrain. To the northwest, the Népoui River flows into another bay where there has been rapid advance of mangroves on an accumulation of red gravelly mining waste brought down by the river from hilltop nickel mines in the hinterland (>Fig. 23.1.6).

Beyond the Presqu’île de Pindaï, the coast is straighter and has no river mouths. Fringing reefs extend along Baie Blanche to the coralline gravel Plage des Franco. At Tarbeville, dark beach sand with shells and gravel forms longshore spits to the north, with mangroves (mainly Rhizophora) in the swales behind them. Near Voh, a dry piedmont slope fronts the steep sided uplands. Presqu’île de Gatope, to the north, has an intricate mangrove shoreline, and the Temala delta (>Fig. 23.1.7) has stunted mangroves  and wide saline flats beside an old wharf with spilled mining material. The shore to the north is rocky, with breccia reefs. Near Ouaco, the Pointe de l’Ounda projects into a sea stained red with mining pollution from the Poué Koué River. Beyond Cape Devard the steep, terraced Iounja River valley crosses the piedmont fronting the Kaala Massif, which has opencast hilltop nickel quarries spilling mining waste down the bordering slopes. Toward Karembé, the piedmont descends to a gravelly beach fronted by dead fringing coral reef, and a cuspate sand ridge curves into a mangrove swamp as a chenier (>Fig. 23.1.8). Progradation of beaches with sand and gravel and tidal flats with red mud are found after heavy rain accompanying Cyclone Gyan in 1981 caused river flooding and the discharge of large ­quantities of mining waste to the shore and nearshore areas. The channel of the Tinip River, aggraded with mining waste in 1980, was sluiced by this flooding to expose formerly buried tree trunks in 1982, and as a result there was progradation of the sandy and muddy shoreline at its mouth.

⊡⊡ Fig. 23.1.5 Steep coast on Léprédour Island, Commune de Boulouparis.

New Caledonia and the Loyalty Islands

23.1

⊡⊡ Fig. 23.1.6 Red-brown mudflats in the Baie de Népoui. (Courtesy Geostudies.)

⊡⊡ Fig. 23.1.7 The “Coeur de Voh” is a part of the Temala Delta, is a former salted flat area in the middle of a Rhizophora forest. By 2002, this heartshaped area had been colonised by 2 m high Avicennia mangroves.

The Iouanga River also carries large amounts of sand and gravel derived from weathering outcrops of phthanite and Pleistocene valley terraces dominated by this material. Its delta is swampy, with cheniers of grey sand, shells and coral deposited on saline flats and corridors of mangrove.

Near Koumac is Pointe Pandop, a reef-fringed coastal hill of conglomerate backed by a lagoon. The Koumac delta to the north is bordered by brown gravelly beaches and spits that have grown northwestward, extending past the cuspate foreland of Mngadiou (in the lee of a patch reef) and on toward Paagoumène, to end in a recurved

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New Caledonia and the Loyalty Islands

⊡⊡ Fig. 23.1.8 Chenier in mangroves at Karembé. (Courtesy Geostudies.)

spit. This grew about 100 m between 1954 and 1976, prograding the coastline by about 250 m. Behind and beyond this spit the shore is mangrove-edged, backed by old beach sands and a broad plain. A cay, Ile Targadiou, stands on a patch reef offshore. North of Paagoumène, the coast steepens, and there are headlands of serpentine between bays with beaches of dark sand and gravel from the piedmont deposits, and sectors of white coralline sand and gravel behind an almost continuous fringing reef. To the north, the River Néhoué flows into the broad Baie de Néhoué, bringing down red clay, sand, and gravel from the hilltop mines in the Tiébaghi Massif. North of Poum, the beaches are dominated by reef-derived calcareous sand and gravel. The barrier reef continues northwestward offshore as the peninsula narrows to Pointe Naharian and the Ile Baaba. To the east are the reefs and islands of the northern tip of New Caledonia in waters sheltered by barrier reefs on either side. The River Diahot, the longest (90 km) in New Caledonia, flows northward down a valley that opens through a tidal estuary to a funnel of swamps and shoals in Baie d’Harcourt. Fluvial deposition is proceeding rapidly to form a swampy deltaic plain studded with ­numerous shallow lakes and saline plains, and this deltaic progradation has cut off an old rocky coastline from the sea toward Arama. On the eastern side, the Massif du Panié declines to a ridge ending at Pam, Ile Balabio being its resumption in waters where coral reefs are very extensive. The beaches are dominated by reef-derived calcareous sand and gravel.

The east coast of New Caledonia is steeper and more exposed to waves generated by the southeasterly trade winds, especially where the barrier reef (which excludes most oceanic wave action) is outlying. The eastern barrier reef, the Récif de Cook, runs in from the northwest and is backed by shoals and the steep coast of the Massif du Panié (>Fig.  23.1.9). This has rocky watercourses, scrubby gulleys, minor inlets at valley mouths, fringing reefs between with minor beaches mainly coralline. There are rocky outcrops on headlands, this is a low wave energy coast because of the outlying reefs and cliffing is limited. Balade is the site where the first French missionaries landed in 1843 to establish a church. A grey sandy beach curves out to Pointe Sainte Matthieu, where the inner barrier reef diverges southeast from the coast, and there are wide mangroves fronting lagoons and salt flats toward the village of Pouébo. The mangrove-fringed coastal plain narrows toward Oubatché, and toward Cap Colnett, the mangroves behind the fringing reef fade out as exposure to wave action increases. The land rises abruptly to Mont Panié (1,628 m), the summit of the Massif du Panié and the highest peak in New Caledonia. At Ouaième, the river flows into a deep inlet, and to the south is a prominent recurved spit of white sand that has drifted northwest along the coast (>Fig. 23.1.10). To the southeast is the Baie de Hienghène, a deep ria where the Tanghène and Hienghène Rivers flow into a shoaly bay behind a coast notable for the bare Lindéralique cliffs, which run along a narrow steep-sided ridge of dark grey  crystalline Eocene limestone, up to 60 m high, with

New Caledonia and the Loyalty Islands

⊡⊡ Fig. 23.1.9 Steep coast, Massif de Panié. (Courtesy Geostudies.)

⊡⊡ Fig. 23.1.10 Recurved spit at Ouaième. (Courtesy Geostudies.)

23.1

jagged rocks worn by karstic weathering. There are caves and coves, and the cliffs show basal notches, cut by solution of the limestone, and as prominent on the inner, sheltered shore as on the outer exposed side. Le Sphinx and Le Poulet are tall tower karst limestone stacks in the bay, and to the southeast a fringing reef borders the Tours Notre Dame and a limestone ridge breached by an inlet to the mangrove-fringed Lagune de Lindéralique. Severe flooding has found here during cyclones, as in 1981, when the rivers rose 17 m during the Cyclone Gyan downpour. The tower karst comes to an end, but fringing reef and coralline sand beaches continue to the southeast, interrupted by the estuary of the Tipindjé, which is bordered by paired spits. Mangroves fringe estuarine inlets, and terrigenous pebbles, cobbles, and boulders diversify the ­coralline beaches at the mouths of gulleys descending the steep coast. At Touho, the outer reef (Récif Mengalia) curves in to the coast, which here swings southward into a bay that curves round to the Poindimié promontory. There are many cays and islets offshore. The Récif de Poindimié shelters this bay, which has beaches shaped by refracted southeasterly waves. The Tiwaka and Amoa Rivers flow down to estuaries with mouths constricted by paired spits of grey sand. Much use has been made of coral limestone quarried from the nearshore area for roads and buildings, as well as breakwaters for village harbours, as at Tibarama. South of the promontory, another bay curves past a series of river mouths. The mangrove-fringed estuary of the Yahoué and Tchamba Rivers south of Ometteux has

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shoals of grey mud, sand, and gravel, and is bordered by a cuspate foreland with parallel beach ridges. The Nabai (Nimbayé) River flows into a wide meandering estuary beside the town of Ponérihouen, bordered by beach ridge plains of fluvially supplied and coralline sand shaped by wave action. South of Mou, a fault at Kora Creek marks the beginning of the peridotite rocks, which form steep-fronted uplands on the east coast of New Caledonia, and the fringing reef is stained by red mud. Several rivers are delivering sand and gravel loads augmented by the effects of mining within their catchments to nourish beaches that are exten­ ded northward by longshore drifting. Monéo River, joined by the Néavin, opens through a swampy delta with mangrove islands fringed seaward by beaches and beach ridges. The beaches contain much black pisolitic sand derived from lateritic crusts on the peridotites within the river catchments, and new spits were built here as a sequel to the flooding that accompanied Cyclone Gyan in 1981. To the south, Hô River is also bordered by beaches of coarse black pisolitic sand. Cape Bocage is a high peridotite peninsula extending southeast, sheltering Baie Lebris, where mangroves are growing on red mud brought down by rivers from the nickel mines in the hinterland and carried northward by longshore drifting. The steep Bâ River cascades down a bouldery channel into the head of this bay. To the south, the coast has a sandy beach backed by beach ridges and interrupted by outlets from the Houailou

estuary, off which there are variable shoals. The Houailou River channel was widened by floods in Cyclone Gyan 1981, and opens into an estuary with mangroves on deltaic islands, where it is joined by the River Néaoua, which beings down a heavy load of mining waste. A coastal ridge develops behind Paraouyé, with a fringing reef and coralline beach that has beach rock exposures. In the Bay of Poro to the southeast, waste from hilltop mines has been discharged down the coastal slope and hillside gulleys to form fans extending on to the shore near a wharf. This has augmented a beach of dark pisolitic ironstone sand and granules, locally cemented to beach rock. Nekoué stands on beach ridges of sand and gravel over a fringing reef, and northward drifting has built a spit at the mouth of Nekoué River. The steep coast continues, interrupted by the Koua River valley, to the Baie de Kouaoua. Here the Kouaoua River, carrying reddish sediment derived from mining waste, opens through a mangrove-fringed estuary to an outlet between paired sand spits. The estuary, mangrove mud, and nearshore sea are discolored by the red mud brought down by the Kouaoua and also the Karoipa (Kakenjou) River, entering from the south. From the mouth of the Kouaoua, a beach of grey sand with some pebbles and shells runs south to a cuspate foreland, and then to the Aoumou spit, which has grown southward in front of a beach ridge plain, Longshore drifting is southward in the shelter of the large peninsulas on the southern side of the Baie de Kouaoua. ⊡⊡ Fig. 23.1.11 Aggraded by mining waste, the Ouano River is building a delta out into the Baie de Canala. (Courtesy Geostudies.)

New Caledonia and the Loyalty Islands

To the south is the long, narrow Baie de Canala, with a high peninsula on its eastern side. The Ouango River, laden with mining waste, flows through a braided valley floor to a growing delta (>Fig.  23.1.11) on the western coast of the Baie de Canala, while the Canala River, similarly laden, flows into a mangrove-fringed ria at the southern end. Farther south, the River Nakéty flows through an estuary with shoals of black sand into the head of another large bay, its mouth bordered by a beach ridge plain that is being eroded by the sea. A straight, steep coast with a rocky shore runs southeast to the mouth of Dothio River, which enters the sea through a beach ridge plain. The Dothio is a clear stream flowing over boulders and gravels until it is joined by a right bank tributary carrying a massive gravelly load spilled from the nickel mines on Le Plateau to the south. The estuary has shoals of mining waste. To the south, the River Thio drains a broad valley and is also laden with gravel and red mud from the hilltop nickel mines (>Fig. 23.1.12), which has filled a formerly deep channel. Sand spits flank the entrance, and an upper beach of black sand above gravelly beach conglomerate resumes south of Bota Méré, a prominent hill that was formerly an island, and runs behind Baie de la Mission into which the River Thio formerly flowed. Considerable changes are found around the mouths of these rivers during the floods that accompanied Cyclone Gyan in 1981, when waves subsequently rebuilt spits and bars of sand and gravel washed into the sea by flood waters. ⊡⊡ Fig. 23.1.12 Mines Bornet, showing mining waste spilling down the hillside and the resulting gravelly load of the River Thio. (Courtesy Geostudies.)

23.1

Southeast of Thio, the coast remains steep, punctuated  by steep-sided valleys such as those of the N’Deu, N’gove, and Ouinné Rivers. The Yaté River, modified by the building of a large barrage upstream, flows into a marine inlet, and along the coast to the southeast a steep slope descends to an apron of downwash that extends out over an emerged Pleistocene fringing reef 25 km long, on which the villages of Touaouru and St. Gabriel stand. There is also a modern fringing reef, and a barrier reef that diverges southward. Port Boisé is a small mangrovefringed bay. The steep coast with a fringing coral reef runs round the southern part of New Caledonia to the large ria of the  Baie de Prony, and on past the outlying Ile Ouen to the southwest coast. Beyond the Baie des Pirogues is the sandy Baie de Mouéa, bordered by the reef-fringed Mont Dore peninsula, rising to 772 m. Small bay barriers enclose mangroves and the fringing reefs are backed by beaches of shelly sand and coral gravel. Offshore are Ile  Ndé and Ile Bailly, both with trailing spits at their northern ends. West of the Baie de Boulari is Port Ngéa, with a ­mangrove-fringed shore behind a fringing reef. A beach  develops and becomes a cuspate foreland below Mount Ouen Toro, then runs behind the wide bay of Anse  Vata round to a rocky headland. To the north is the  Baie  des Citrons with a sandy beach, then the Baie de  l’Orphelinat and the promontory on which Nouméa stands.

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New Caledonia and the Loyalty Islands

3.  Isle of Pines The Ile des Pins, 50 km southeast of New Caledonia, consist of a central plateau of laterised peridotite, rising to Pic Nga (262 m), surrounded by emerged Pleistocene coral reefs. Their restricted extent and varying altitude indicate that tectonic uplift has taken place here. The emerged reefs end seaward in cliffs, which have basal notches and visors produced by solution (>Fig. 23.1.13) and some small segments of shore platform reduced to just above low tide level, also largely by solution processes. These in turn are fringed by extensive Holocene coral reefs, the Récif Tiaré, built up to this same level. These reefs permit attenuated swell and waves generated by the southeasterly trade winds to reach the island coast, shaping beaches, notably in the Baie de Kuto and the Baie de Kanumera, on either side of the isthmus (tombolo) that links the former island of Kuto in the southwest. Near Vao, the sandy beach is backed by a foredune 2 m high, backed by beach ridges. There are cays of sand and coralline gravel, such as Ilot Brosse, out on the Holocene coral reefs. Beach rock is exposed locally where calcareous sandy beaches have been eroded.

4.  The Loyalty Islands The Loyalty Islands are a chain of emerged reef formations parallel to, and northeast from, New Caledonia. The

southernmost, Walpole Island, is a patch reef raised 75 m above present sea level. Maré has the form of a raised atoll with a rim 60–138 m high, the southwest being higher than the northeast. It is surrounded by a modern fringing reef, and stages in elevation are marked by 15 terraces, the upper ones tilted, those below 12 m horizontal. In the centre are three basaltic peaks, all that can be seen of the basement, a volcanic island of early Miocene age around which the reefs developed and were afterward uplifted. Lifou is similar, but the raised reef rim reaches only 104 m above present sea level. Five terraces mark stages in its uplift. Ouvéa has the form of a tilted atoll, the eastern rim raised up to 46 m above the present sea level, with terraces on its seaward flank (>Fig.  23.1.14), the western rim submerged, but with a chain of younger reef segments built up from it to enclose a lagoon with a floor that declines westward. Farther north, BeautempsBeaupré is similar, the southeastern part raised only 4 m above low-tide level, while the Astrolabe Reefs beyond are not emerged features. Although it has been generally believed that Maré, Lifou, and Ouvéa were atoll formations raised by tectonic movements and since modified by karstic wea­ thering, it has been alternatively suggested that they originated from reef platforms that suffered differential biochemical erosion after uplift, with excavation of the central areas and persistence only of harder rims of dolomitised rock.

⊡⊡ Fig. 23.1.13 Cliff on emerged coral limestone, Baie de Kanumera, Ile des Pins, showing basal notch. (Courtesy Geostudies.)

New Caledonia and the Loyalty Islands

23.1

⊡⊡ Fig. 23.1.14 Terraces mark stages in the emergence of the east coast of Ouvéa. (Courtesy Geostudies.)

References Bird ECF, Dubois JP, Iltis JA (1984) The impacts of opencast mining on the rivers and coasts of New Caledonia. United Nations University, Tokyo, Japan

Chevalier JP (1973) Coral reefs of New Caledonia. In: Jones OA, Endean R (eds) Biology and geology of coral reefs. Academic Press, New York, pp 143–167 Stevenson J, Dodson JR, Prosser IP (2001) A late quaternary record of environmental change and human impact from New Caledonia. Palaeogeogr Palaeoclimatol Palaeoecol 168:97–123

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23.2 Fiji

Patrick Nunn

1.  Introduction The Fiji Islands are composed of a group of some 350 islands (around 90 inhabited) in the tropical southwest Pacific. The character of particular coasts depends on several factors, including aspect, island geology and lithology, island size (particularly whether large rivers are present or not), island tectonics, and offshore coralreef configuration (important for the supply of calcareous sediment). Add to these natural factors a highly variable degree of human impact, ranging from the negligible on sparsely inhabited islands to the severe along coasts which have been intensely developed for tourism, for example, and you find an almost bewildering diversity of coasts and coastal dynamics which make coastal management very challenging. This chapter looks at each of the above-mentioned factors, with examples given, before going on to a systematic account of the Fiji coastline.

2.  Aspect All parts of the Fiji Islands are affected by the southeast trade winds, which are strongest during the winter months (June–September). Windward coasts, therefore, commonly have more serious coastal erosion problems, although this is offset on many by the presence of broad offshore reefs, which develop to their fullest extent off the windward sides of islands. A good example is provided by the island Ovalau, along with its neighbour, Moturiki. The broadest reefs have formed along the windward coasts, whereas along the leeward coasts, reefs are narrower and less continuous. Also, villages such as Nasinu along the windward coasts of Ovalau have more severe erosion than those along the leeward side. Leeward coasts are not immune from coastal erosion. In Fiji, most tropical cyclones (hurricanes or typhoons) approach from the northwest, so that many coasts in these locations exhibit the effects of these storms. These effects vary depending on the antecedent conditions. For example, in a location where there is little ocean-floor sediment off the coast, the

effects of a tropical cyclone may cause massive amounts of erosion: a situation in Fiji which is typified by the exposure of beach rock ramparts. Alternatively, where such sediment are abundant, storm surges associated with tropical cyclones may dredge up such material and deposit it along the coast resulting in long-term coastline progradation rather than retreat. Transitional coasts in Fiji are those that are neither strictly windward nor leeward. Often they are dominated by processes associated with terrestrial processes, although it takes only one tropical cyclone to impact the coast at right angles for the signature of these terrestrial processes to be obscured. A good example of a transitional coast is that of southwest Ovalau Island, where sediment from the Bureta River has led to coastal progradation and, as shown below, to mangroves advancing over coral reef smothered by fluvial sediment. Aspect is probably the single most important criterion in determining the coastal character in Fiji, but it is rarely the sole one.

3.  Geology and Lithology The main group of the Fiji Islands, comprising Vitilevu and Vanualevu and nearby islands (including those in the Koro Sea, Kadavu, and the Yasayasamoala) are of largely igneous origin. The larger islands are amalgams of former island arcs and display a variety of igneous, metamorphic, and sedimentary rocks, particularly in their interior parts. The coastal fringes of these islands are composed largely of sediment derived from erosion of the inland massifs; bedrock is commonly exposed only on headlands. The older igneous rocks typically form gently rolling terrain which meets the coast less abruptly than the younger rocks which are often cut into steep cliffs along the coast. In contrast, the sediment-composed coasts of the main islands of Fiji accumulated mostly only since the fall of the sea level about 4,500 years ago, and are sites of coastal progradation. There are often series of beach ridges marking the gradual fall of late Holocene sea level and the accompanying progradation.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_23.2, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Fiji

When people arrived first in the Fiji Islands about 3,000 years ago, they occupied the landward sides of these beach ridges. Finding their settlements today has aided the understanding of late Holocene coastal development in Fiji. The Yasawa and Mamanuca groups of western Fiji are mostly (in terms of surface area) of igneous composition, but also include many reef islands – cays, motu, and emerged reef islands in the sense used by Nunn (1994). Most of the cays and motu are mobile over decadal time scales, and this is proving a problem for some of those that have been developed for tourism and require a degree of stability. A good example is provided by picturesque Eluvuka (Treasure Island) (>Fig. 23.2.1). Monitoring of the coast of Eluvuka has shown that sand moves around the island depending on dominant wind direction (Nunn 2001). Here, in the Mamanuca Lagoon, the islands are largely shielded by Vitilevu from the southeast trade winds; particularly, during El Niño events, the wind direction tends to be mostly from the northwest. Sand moves around these islands away from the dominant wind and wave direction. On Eluvuka, beach rock has been exposed on the western shore as the result of sand loss. The Lau Islands of the eastern Fiji are distinct from the rest of Fiji because of their dominant limestone composition. Most such islands have cliffed coasts with beaches common only in structural (rather than erosional)

e­ mbayments. The example of the southern coast of Nayau Island, exhibiting at this location, a series of three limestone terraces (>Fig. 23.2.2).

4.  Island Size The larger islands of Fiji generally have a greater diversity of coastal types than the smaller islands. This is not only because of the climate and geology, as discussed earlier, but also because terrestrial processes, particularly fluvial erosion and deposition, are commonly more vigorous on larger islands. The coast of the largest island, Vitilevu, is marked by huge deltas of fluvial and marine sediment, which provide environments for intense mangrove ­development, for example. Delta coasts in Fiji are also those which change most rapidly. For example, the Sigatoka Delta in southwest Vitilevu is a wave-controlled feature, the seaward extremity of which is characterised by an area of sand dunes (>Fig.  23.2.3). Research has shown that the main dunes here accumulated only within the last 800 years when coastal people moved up the valley in ­numbers large enough to impact the vegetation and help release huge quantities of terrigenous sediment into the river. Carried to the sea by the river, this sediment was then driven onshore by waves from the southeast and thence blown inland by the southeast trade winds to form the dunes.

⊡⊡ Fig. 23.2.1 View of Eluvuka (Treasure Island), a typical sand cay in the Mamanuca Group. (Courtesy Treasure Island Resort.)

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⊡⊡ Fig. 23.2.2 The south coast of Nayau Island, central Lau Group, showing three reef-limestone terraces which mark stages in the emergence of this part of the Lau Ridge.

⊡⊡ Fig. 23.2.3 The Sigatoka sand dunes formed mostly only within the last 700 years as a result of increasing suspended sediment loads in the Sigatoka River, perhaps a consequence of increased forest clearance by human settlers upstream.

5.  Tectonics Most of the Fiji islands lie on a small lithospheric plate – the Fiji Plate – which is one of several microplates between the giant Pacific Plate and Indo-Australian Plate. These two huge plates converge obliquely in the Fiji region with the result that for the past 40 million years or so, the Fiji

Platform has been rotated anticlockwise. Continuing stress on the Fiji Platform is manifested in the rise and fall of various islands (Rodda 1994). The Lau Islands are somewhat different, in that the ridge from which they all rise has been emerging for the past 5 million years or so as a result of the heating-up of a detached, partly-subducted slab of the Pacific Plate

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⊡⊡ Fig. 23.2.4 The central rift valley of Moala Island (seen here from left to right) is partly filled by Uciwai Bay (on the left) with Keteira Village at its head.

⊡⊡ Fig. 23.2.5 View of Gusuniqara Point, northwest Vatulele Island, showing the different levels of emerged coastline notches. Each notch is believed to have been uplifted from sea level by a coseismic uplift event.

lithosphere. The island Cikobia (Macuata) in northeast Fiji is located close to the transverse plate boundary named the Fiji Fracture Zone and is rising rapidly for that reason (Nunn 1994). Most subsiding coasts in Fiji can be identified by their embayed character. The stellate (star-shaped) island of Ono in the Kadavu group was hailed as a classic example by Davis (1928). In many parts of the Yasawa Islands, ­similar embayments can be recognized. Recent research

also identified the north coast of Vitilevu Island as having been subsiding since the middle Holocene at a rate of 0.2 mm/year. There are signs of submergence all along this coast. Even more recent research in the northern Lau Islands found that while most parts of the area have been stable during the latest Quaternary, the Vanuabalavu Island group was in fact subsiding, at a rate of around 0.25 mm/ year (Nunn et al. 2002). This explains why the coastal

Fiji

⊡⊡ Fig. 23.2.6 Mushroom rock 35–40 m above sea level on Lawarea Point with a person beneath it for scale. The notch at the foot of the rock is thought to have been cut by the sea during the late Quaternary before the land was uplifted.

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Uplift has also affected parts of the larger Fiji Islands. The main island of the Kadavu group, for example, is divided into three mesotectonic blocks, the westernmost of which has risen relative to the others. Only around this western block are there fragments of emerged reef; the island Dranubu with the large young volcano named Nabukelevu (Mt. Washington) in the background. The south coast of Vitilevu Island has also risen in places, probably because it adjoins an active seismic zone offshore. Furthermore, there is evidence from emerged reefs at Korotogo near Sigatoka. The southeast part of Vanualevu Island is composed of the Cakaudrove Peninsula and exhibits along its southern parts much evidence for co-seismic uplift. Similar evidence for co-seismic uplift occurs on Vatulele Island in the southwest Fiji (>Fig. 23.2.5). Uplift is also evident along much of the north coast of Vanualevu (>Fig. 23.2.6), particularly, its eastern part and nearshore islands. This uplift is believed to be aseismic and associated with movements along the transverse Fiji Fracture Zone to the north.

6.  Coral Reefs

erosion attributed to recent sea level rise has been so much more severe in this island group than elsewhere in the northern Lau Islands. Elsewhere in Lau, late Quaternary uplift has occurred, some in association with point loading of the ocean-floor lithosphere by young volcanoes such as Moce and Kabara. The Yasayasamoala (Moala Island group) is a group of three islands and one atoll in the southeast of Fiji which have partly ascended a flexure in the ocean-floor lithosphere. Parts of Matuku Island have therefore risen several tens of metres, while there are no such indicators on Totoya Island. The largest island in the group – Moala – has experienced rifting as a consequence of rising up the flank of this flexure, and this has produced rapid subsidence in Uciwai Bay (>Fig. 23.2.4).

The configuration of offshore coral reefs in the Fiji Islands is complex and reflects the Quaternary history of fluctuating sea levels, tectonics, and varying oceanographic conditions in space and time. The largest ring reefs are found in the northeast and east parts of the group with many atolls and almost all atolls found in these areas: an example is Wailagilala Island. In the west of the group, the Great Sea Reef (na Cakau Levu) encloses a large lagoon in which the Mamanuca and Yasawa Islands are located. This reef extends off the north coast of Vanualevu. Many patch reefs are found within this lagoon which boasts many fine island beaches and is one reason why so much of Fiji’s tourist development is located there. In contrast, the reefs which surround most of the larger islands are less substantial in both breadth and configuration. The famed Coral Coast of southwest and south Vitilevu Island has actually only quite a narrow fringing reef, which is why it is sometimes affected by tsunami and storm surges when other coasts are spared. Many nearshore reefs, in particular, have been degraded as a result of terrestrial sedimentation, pollution, and direct human impact. This affects their ability to supply calcareous sediment to adjacent beaches, and some

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Fiji

⊡⊡ Fig. 23.2.7 Map of Fiji showing the main islands mentioned in the text. 180°

178°E vu

ale

sa wa s

17°S

nu Va

Ya

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Koro Ovalau

Vitilevu

Taveuni Vanua Balavu Mago Cicia Lakeba

Gau Vatulele 19°S

Beqa

50 km Kadavu

Moala Matuku

Moce Kabara

Totoya

recent beach erosion has been attributed to such a disruption to sediment supply (Nunn 2001).

7.  The Coastlines of Fiji The Fiji Islands fall into several distinct groups (>Fig.  23.2.7). The two largest islands, Vitilevu and Vanualevu, will be considered separately. The islands of central Fiji (Lomaiviti) including Ovalau and Moturiki, and Taveuni are considered in another group. The Yasawa and Mamanuca groups are discussed together. The Kadavu group, the Yasaya-samoala (Moala group), and the Lau group are each considered separately. No ­specific ­infor­mation is available on distant outliers, Rotuma and Ceva-i-ra.

7.1.  Vitilevu The coast of Vitilevu island is dominated by large deltas, mostly wave-controlled. They include the Rewa, the Navua, the Sigatoka, the Nadi, the Ba, and the Nasivi (Tavua) deltas. Most of these deltas are active and covered with mangrove forests (>Fig. 23.2.8). At the fronts of the wave-controlled deltas, erosion generally dominates and excess fluvially-transported sediment moves offshore. Only in parts of the deltas of the Rewa and Ba there is significant progradation. In the case of the Rewa, this is because the river sediment load is so great that even waves driven by the southeast tradewinds cannot prevent progradation. In the case of the Ba, although river sediment load is also large, the river empties onto a shallow leeward shelf where wave energy is comparatively subdued.

One of the main sites of interest on the Vitilevu coast is the Sigatoka River mouth, where the large dune field has attracted the attention of many geomorphologists over the years and, because of the early human remains which are associated with it, has attracted the interest of archaeologists. Within the dunes, there are three palaeosols (ancient soils) in which cultural remains of distinct ages are found, the oldest dating from the time about 2,800 years ago of the earliest known humans in Fiji, the so-called Lapita people. The presence of these palaeosols within an otherwise archaeologically-sterile dune sequence poses many questions about Holocene environmental change. It seems most likely that changes in climate, in sea level, river sediment loads, and perhaps even delta channel networks, were responsible for creating environments where soil supporting forests existed in this area from time to time during the past 3,000 years. The main dunes appear to have accumulated only within the past 800 years when people first moved in large numbers inland along the fertile Sigatoka Valley (>Fig. 23.2.9). Another area of interest is a little to the east of the dunes where a fossil coral reef is exposed at the back of the modern beach (>Fig. 23.2.10). Repeated attempts to date this reef have proved futile because of the comprehensive recrystallization of the corals within it. It may be a Holocene feature, in which case, it is certainly been uplifted as well as exposed by sea level fall, or it may be an earlier Quaternary reef (perhaps Last Interglacial). Farther to the east, along the western borders of the Navua delta plain, a series of Holocene beach ridges have been mapped and sampled. This work shows that a Holocene embayment was closed by a barrier beach during Holocene sea level fall and how a series of beach ridges formed subsequently, while at the same time, the backlagoon gradually infilled (Shepherd 1988, 1990). The backlagoon area is now occupied by the Pacific Harbour complex. To the east of Suva, and south of Laucala Bay, the islands of Nukulau and Makuluva are sandy cays perched on coral reefs. Makuluva Island has migrated westward, as shown by its eroded eastern shores with offshore remains of wells formerly drilled through the cay down to the reef, and prograding western shores.

7.2.  Vanualevu Large deltas are fewer in Vanualevu than Vitilevu, although the merging of three rivers in the Labasa area has given rise to a large delta, clearly wave-controlled at the mouth

Fiji

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⊡⊡ Fig. 23.2.8 Near the edge of the Navua Delta, south coast Viti Levu. The village of Vunibau on the inland side (right side of photo) lies within 100 m of the delta edge, close to the mouth of the main river. Some 25 years ago, a bank of sand and gravel was thrown up at the front of the delta by a storm surge. This created a narrow strip of land (left side of photo), now covered with mangrove forest, which now protects Vunibau from coastline erosion.

⊡⊡ Fig. 23.2.9 The lower Sigatoka River valley. Like many of the large rivers on Viti Levu and Vanua Levu islands in Fiji, the suspended sediment load of this river significantly alters the nature of the coasts that adjoin its mouth.

of the Qawa (>Fig. 23.2.11). The large delta of the Dreketi River, the largest on Vanualevu, is on the leeward coast of the island and is consequently not wave-controlled but is actively growing seaward. Waves trim prograding delta margins at the mouths of south-coast rivers, the Yanawai and Wailevu, which empty into a large shallow embayment along the south coast.

The southwest coast of Vanualevu is dominated by landforms associated with the huge intraplate volcano of Seatura, which has become accreted onto the distinct island-arc volcanoes which form the rest of the island. The southeast coast (the Cakaudrove Peninsula) is dominated along its windward side by uplift, and there is evidence here of coseismic uplift (>Fig. 23.2.5). To the north, the

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⊡⊡ Fig. 23.2.10 The emerged reef at Makasiko, just east of the mouth of the Sigatoka River.

⊡⊡ Fig. 23.2.11 Aerial view of the mangrove-fringed Qawa River mouth and adjacent north coast of Vanua Levu.

large rift which almost bisects the island is occupied by Natewa Bay. There may be some subsidence along the northern coasts of this bay which are exacerbating the effects of recent sea level rise here.

8. Islands of Central Fiji (Including Ovalau and Moturiki, and Taveuni) The islands of Ovalau and Moturiki, off the east coast of Vitilevu, are thought to be one of the most tectonicallystable parts of Fiji, for which reason, they have been studied intensively in the expectation that they would reveal the effects of sea level changes alone. Work on Holocene sea level changes around Ovalau and Moturiki revealed two peaks of sea level separated by a prominent gap or hiatus (Nunn 2000).

As elsewhere in Fiji, most of the data used to reconstruct the course of Holocene sea level changes came from radiocarbon dates on emerged microatolls of the species Porites lobata. The hiatus in the sea level data from these islands is thought to be significant rather than an artifact of data collection, and may indicate early human impact, perhaps through the inadvertent introduction of an exotic predator, on reef-surface corals. In the case of the Ovalau-Moturiki hiatus, there is only one contemporary indicator of environmental change which could be linked to human impact (at Bonatoa Bog), but all material evidence of human arrival is much later (Nunn 2001). This evidence is shown for the arrival on nearby Naigani Island and for Ugaga (Stewart) Island in the Beqa lagoon. Besides providing an important benchmark for Holocene sea level changes, the Ovalau-Moturiki area has also

Fiji

been of interest to those interested in evaluating the effects of twentieth-century sea level rise on Pacific Island coastlines without interference from tectonics. The effects of coastline erosion around these islands are manifested most clearly on windward coasts from which the mangrove fringe has been cleared. It is clear that windward coasts are more vulnerable to erosion, particularly those which face wide reef gaps (called daveta in Fijian); the Ovalau villages of Tokou and Nasinu are good examples. Erosion along the front of Nasinu is shown in (>Fig. 23.2.3). Erosion also occurs along the windward coast of Moturiki Island, as seen in this view, north of Daku Village. Yet along the front of Daku Village the sensible positioning of a groyne in the 1970s has led to the progradation of the coastline along the village front by some 30 m. In contrast, the arbitrary positioning of a seawall in Nasauvuki Village on Moturiki has done little to arrest coastline erosion (stimulated by clearance of a belt of mangroves some 600 m broad) or to stop the daily inundation of the village front at high tide. Navutu village on Moturiki is an example of a low-lying village that cannot be set back because it stands at the foot of a cliff. Mangrove clearance has been a major cause of accelerated coastal erosion around these islands, as indeed it has elsewhere in Fiji and the tropical Pacific Islands. This is shown well by two villages on Ovalau – Bureta and

⊡⊡ Fig. 23.2.12 Aerial view of the island Totoya in southeast Fiji. This Pliocene volcano has passed through a caldera stage of formation. The caldera was subsequently breached and drowned. The modern island composed of the flanks of the original volcano. (Courtesy Randy Thaman.)

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Visoto – whose inhabitants have traditional taboos preventing the cutting-down on mangroves. The mangrove fringe around both the villages is well preserved, and the signs of erosion are few, and the effects of storm surges, typically during tropical cyclones, are comparatively muted. Similar problems exist elsewhere in the central islands of Fiji, but they have not generally been studied systematically. The islands most vulnerable to coastal erosion are those young volcanic islands with straight coasts and often only with superficial offshore reef development like Koro and Taveuni. The narrow coastal strips on these islands are often being severely eroded and many settlements are migrating slowly inland in response. On the older volcanic islands with deeply embayed coasts, there is generally less erosion, not only because of the degree of embayment, but also because of the mangroves that commonly line the heads of these bays. Low population densities (except locally) mean that few of these mangrove forests have been impacted by humans as much as those on more populous islands like Moturiki. One interesting, yet barely studied, phenomenon in the islands of central Fiji is the presence of small limestone islands and/or reefs to the south of many of the large volcanic islands. Most islands in the group like Nairai and Gau have such islands. Tiny Yaciwa Island off the southern tip of Gau is an example of such an island.

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⊡⊡ Fig. 23.2.13 The large volcano Nabukelevu (Mt. Washington) at the western end of Kadavu Island was once thought to have last erupted about 50,000 years ago. Yet, with the extension of a road around its base in 1999, a pottery-bearing soil buried by several metres of volcanic ash was discovered, demonstrating that the last eruption must have been within the period of Fiji’s human occupation, that is, the last 3,000 years. Note that the smoke rising in this photo is only from a farmer’s field.

9.  Islands of the Yasayasa Moala

10.  Islands of the Kadavu Group

The Yasayasa Moala (Moala Island group) is a group of oceanic central volcanoes in the southeast of Fiji that appear to have been partly disturbed by movements across a flexure in the oceanic lithosphere. There are three main islands in the group. The easternmost is the horseshoeshaped island Totoya (>Fig. 23.2.12), which has not risen up the flexure and shows no signs of emergence other than those attributable to late Holocene sea level fall. The largest island in the group is Moala, which has risen up the flexure, exhibits a variety of emerged coastline indicators, and has a large central rift valley, marked by Uciwai Bay (>Fig.  23.2.4), which is believed to have resulted from its rise up the flexure. The other island in the group – Matuku – exhibits pillow lavas several tens of metres above sea level, and this has been interpreted as a sign of uplift associated with the island’s journey up to the crest of this now-inactive lithospheric flexure. The fact that the flexure is now inactive – because the convergent plate boundary to the northwest is too – means that the islands of the Yasayasa Moala are currently subsiding, a fact which aggravates the erosional effects of recent sea level rise. So severe have these effects been around Moala, the most populous island, that some coastal settlements have responded in ways which may become more common within Fiji and the highisland Pacific during the next few decades as sea level rise accelerates.

Kadavu and the smaller offshore islands are mostly volcanic, with narrow coastal plains, often confined to the heads of bays and often fringed with mangrove. The main island is dominated by the young volcano Nabukelevu (Mount Washington: 13), which is now known to have erupted within the period of the island’s human settlement, that is, the last 3,000 years at most (>Fig. 23.2.13). The islands of the Kadavu group are divided into three ‘mesotectonic’ blocks. The westernmost of these, on which Nabukelevu is located, has emerged relative to the other two. Fragments of emerged reef dated to the Penultimate Inter­ glacial are found off the coast of this block (>Fig. 23.2.14). The central mesotectonic block is relatively stable, although still exhibits signs of coastal erosion attributable largely to recent sea level rise (>Fig.  23.2.15), while the eastern mesotectonic block, which includes the offshore islands within the Astrolabe Reef, is relatively submergent.

11.  Islands of the Lau Ridge The islands of eastern Fiji mostly rise from the north– south  trending Lau-Colville Ridge, a remnant island arc  which split from the Tonga Ridge about five million years ago. The islands are generally more numerous, closer together, and larger in the northern part of the group because this area has been uplifted more than the south.

Fiji

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⊡⊡ Fig. 23.2.14 The ocean coast of Nagigia (Denham Island), a small island off the westernmost point of Kadavu Island which is made from Penultimate Interglacial reef limestone.

⊡⊡ Fig. 23.2.15 Coastal erosion along the leeward coast of Kadavu Island near Vunisea is attributable to twentiethcentury sea level rise and perhaps the effects of tropical cyclones (hurricanes) of increasing frequency and magnitude within the past 40–50 years.

Island geology shows that uplift began during the late Pliocene and is continuing (Ladd and Hoffmeister 1945). Many of the islands (like Tuvuca and Vatoa) are wholly limestone; Vatuvara has a limestone coast (>Fig. 23.2.16). Others have their volcanic foundations exposed in a narrow strip along the coast (like Nayau), while some are termed makatea-type islands, having a volcanic centre with an encircling fringe of emerged reef limestone (like Lakeba).

There are a few young volcanoes within the Lau group, the origin and significance of which is not clearly understood because they postdate the split from the Tonga Ridge which led to the abandonment of the Lau Ridge as a site of island-arc volcanism. One of these volcanoes is Moce Island. It has been suggested that the evidence for coseismic uplift on islands surrounding these volcanoes is consistent with the effects of their point-loading of the lithosphere.

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⊡⊡ Fig. 23.2.16 The island Moce in central Lau is an anomalously young volcano which may have acted as a point load on the lithosphere causing uplift of islands nearby.

⊡⊡ Fig. 23.2.17 The east coast of Mago Island, northern Lau. The modern reef-bounded lagoon is to the left, an emerged analog is in the centre, and an even more emerged one to the right. This is one of the most unusual and spectacular pieces of evidence for emergence along the Mago coast.

Much work has been carried out on the north of the Lau group, particularly in the Mago-Vanuabalavu area. Mago is a classic makatea-type island, with a spectacular series of emerged lagoons along the northeast coast (>Fig. 23.2.17). Its tectonic history appears to have been similar to that of most of the rest of northern Lau (excluding the Vanuabalavu group) which involves general uplift during the Quaternary with probable stability (or very slow uplift) during the Holocene. On the nearby islands of Yacata, Kaibu, and Vatuvara, there are a series of emerged coastlines which are thought to have been cut during various late Quaternary interglacial high sea level stands. Examples of two from Kaibu are shown (>Fig. 23.2.18). The similarity in level between at least the Last Interglacial coastline on these islands (5.1–5.3 m above

the present coastline) and the estimated Last Intergla-cial sea level from places like the Huon Peninsula in Papua New Guinea (+6 m) LINK suggests that this area has been stable during the latest Quaternary. In contrast, the Vanuabalavu group appears to have subsided during this period. Indicators of the Holocene sea level maximum are lower than elsewhere and coastal erosion attributable to recent sea level rise is conspicuously greater (Nunn et al. 2002). Although sea level changes within the last millen­ nium  are notoriously difficult to detect because of the problems of dating palaeo-sea level indicators accurately, there is one site on Vanuabalavu where a higher-thanpresent sea level about 700 years ago has been detected. The site is Qaranilaca (cave) at the southernmost tip of the island, where a shell conglomerate has been dated to

Fiji

23. 2

⊡⊡ Fig. 23.2.18 The 5.1–5.2 m emerged notch at Dakui along the east coast of Kaibu Island is believed to have been cut during the Last Interglacial, about 125,000 years ago.

ad 1280–1420. The discovery of this site confirms earlier ideas about a Pacific-wide sea level fall around this time (Nunn 2000).

12. Islands of the Yasawa and Mamanuca Groups (Western Fiji) The Yasawa group is made of largely elongate high volcanic islands of complex structure. The limestone islands of Sawailau – famous for its cave – and outlying Viwa are exceptional. The volcanic islands tend to be deeply embayed with mangrove fringes in places but the absence of large rivers tends to mitigate against the spread of mangroves. Viwa is of interest because it is one of the few places in Fiji where dates have been obtained for Last Interglacial limestones. The Mamanuca group also includes some high volcanic islands (like Malolo), but is composed mostly of low sand cays, many of which are developed for tourism (>Fig. 23.2.1).

13.  Conclusion The coasts of Fiji thus show exceptional diversity in lithology, origins, tectonic histories, and current condition. It needs exceptional managers to ensure that the essential characters of these coasts are sustained, not only for their instructive value, but also because most of the people who occupy them can continue to use them to support their livelihoods. These coasts are not only home to a range of

medicinal and otherwise useful plants, but also host ecosystems which provide a range of foods to coastal dwellers. Effective management is even more of a priority for the coasts of Fiji because so many of them are being eroded and/or inundated with unprecedented rapidity causing social disruption. Should the sea level rise estimates of the IPCC prove correct, then the amount of future disruption will increase significantly.

References Davis WM (1928) The coral reef problem. American Geographical Society, Washington, DC (Special Publication 9) Ladd HS, Hoffmeister JE (1945) Geology of Lau, Fiji. Bernice P. Bishop Museum [Honolulu], Bulletin 181 Nunn PD (1994) Oceanic islands. Blackwell, Oxford, 418 pp Nunn PD (2000) Significance of emerged holocene corals around Ovalau and Moturiki islands, Fiji, southwest Pacific. Mar Geol 163:345–351 Nunn PD (2001) Ecological crises or marginal disruptions: the effects of the first humans on Pacific Islands. N Z Geogr 57:11–20 Nunn PD, Ollier CD, Hope GS, Rodda P, Omura A, Peltier WR (2002) Late Quaternary sea level and tectonic changes in northeast Fiji. Mar Geol 187:299–311 Rodda P (1994) Geology of Fiji. In: Stevenson AJ, Herzer RH, Ballance PF (eds) Geology and submarine resources of the Tonga-Lau-Fiji Region, Vol 8. SOPAC Secretariat, SOPAC Technical Bulletin, Suva, Fiji, pp 131–151 Shepherd MJ (1988) The higher-energy coasts of southern Viti Levu, Fiji, with particular reference to the geomorphology of the Deuba coast. J P Stud 14:1–19 Shepherd MJ (1990) The evolution of a moderate energy coast in Holocene time, Pacific Harbour, Viti Levu, Fiji. N Z J Geol Geophys 33:547–556

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23.3 Society Islands

André Guilcher

1.  Introduction As W. M. Davis (1928) wrote, “no other archipelago in the Pacific coral seas includes so systematic a sequence of island forms as those of the Society group. The sequence begins with a young and reefless volcanic island on the east and ends in several small atolls on the northwest.” This pattern is related to the increasing age of volcanoes from east to west and an associated development of reef growth. According to recent views on ocean floor evolution, such a pattern – which is by no means unique in the Central and South Pacific – may result, as in the case of the Galapagos rise from the migration of the oceanic plate over a hot spot, in the expression of a narrow plume of subsurface mantle material rising and spreading outward. As the plate moves, successive volcanoes come into existence and subsequently become extinct. Fringing reefs develop, and as the islands subside these become barrier reefs and eventually atolls, such as Tizpai, Mopelia (Guilcher et al. 1969), Scilly, Bellingshausen, and Tetiaroa to the west of the Society Islands. Meetia is an uninhabited steep rock 435 m high, bearing only discontinuous fringing reefs, and on Maiao, 154 m high, the expected lagoon is replaced by lakes and fringing reefs.

2.  Tahiti This famous island consists of twin volcanoes, 2,241 m and 1,332 m high, connected by the low Taravao isthmus. The volcanic mountains are deeply dissected by valleys, two of which widen in their upper reaches inside the larger volcano in incipient erosion calderas, and carry appreciable amounts of volcanic pebbles and sand to the sea. Davis paid much attention to the fact that the ridges sloping from the summits to the coasts have been truncated at their outer ends by circumferential cliffs, still visible in the topography. This feature is particularly represented in the southeast. It dates from a time when reefs had not yet been built by the corals around the twin volcanoes. These reefs are now generally barrier reefs, separated from the island by a narrow lagoon, 500–2,000 m in width,

but along parts of the west coast the lagoon is absent and the reef becomes a fringing reef. Where the lagoon exists it commonly reaches 30–50 m in depth. In addition, the barrier reef is drowned in some places, especially in the north, the east, and the southeast (>Fig. 23.3.1), under several metres of water, possibly because of recent local subsidence. In these areas the swell generated by the trade winds penetrates into the lagoon and breaks on the inner shore, so that the pebbles supplied by the rivers become well rounded by the waves, while elsewhere the beaches are sandy and quiet. Numerous passages break the barrier and allow large ships to enter. As in other lagoons behind barriers, some of the sediment have been supplied by local marine organisms, and others by rivers flowing from the central island. Analyses of the calcium carbonate content of these sediment has shown that the influence of the central island is generally large (Guilcher et al. 1969). The carbonate content generally fell below 5% close to the land, and was found to be much smaller around the main volcano (21%) than around the smaller one (66%), because the larger volcano has larger rivers that have supplied more terrigenous sediment to the lagoon. The terrigenous or organogenic nature of the lagoon sediment depends on the ratio between the area covered by the central island and the area of the lagoon.

3.  Moorea Moorea, which is an extinct volcano, is 1,200 m high and has more dissected land than Tahiti. It makes one of the finest skylines of the archipelago, and is surrounded by a lagoon even narrower and more discontinuous than that around Tahiti. The barrier reef, which is wide, merges into a fringing reef in a number of places. There are many ex­­amples of the relationship between passages through the reef and the alignments of drowned valleys, now forming Cook and Papetoai bays. It is evident that the coastline and reef morphology are the outcome of marine submergence.

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_23.3, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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⊡⊡ Fig. 23.3.1 The Society Islands. 1. Barrier reef (sometimes fringing); 2. Drowned barrier reef; 3. Sedimentary apron carried into the lagoon by overwash; 4. Reef alveoles between Tahaa and Raiatea; 5. Tops of volcanoes, with altitudes in metres. (Courtesy Geostudies.)

4.  Huahine Huahine is also considerably dissected, and is divided into two parts by a central channel in which the mean depths reach 30–40 m. It seems that, as on Banks Peninsula in New Zealand, this feature results from the drowning of an eroded caldera. As around Tahiti and Moorea, the barrier reef lies close to the mainland. The lagoon, 20–45 m deep, is discontinuous, and becomes a lake in the north. It is connected to the ocean by five passages, four of them more than 20 m deep. As on many other high islands (notably Bora-Bora and Moorea), fringing reefs have commonly formed on the inner side of the lagoon.

barrier reef in the shape of a figure 8. The lagoon is somewhat wider than in the preceding cases, and where it expands between the two islands coral reefs have grown in the form of large juxtaposed alveoles, a rare but not unique pattern found on a larger scale on Mataiva Atoll in the Tuamotus, Canton Atoll in the Phoenix group, and Christmas Atoll in the Line Islands. In many instances, the passages connecting the lagoon with the open sea bear sand cays (islets) on both sides, a feature by no means unique to Raiatea and Tahaa, but well exemplified here. They have been shaped by swell refraction and diffraction through the passage. An apron of overwashed sediment is seen on the northern side. The relation of passages to inner valleys is as evident as at Moorea.

5.  Raiatea and Tahaa

6.  Bora-Bora

These two islands consist of twin volcanoes separated by a shallow drowned area, but both are surrounded by a single

Many scientists consider that this island displays the most beautiful marine landscape in the world, and Darwin cited

  23.3

Society Islands

it as one of the best examples of a barrier reef island. Because of its western position, the volcano is smaller and the lagoon wider than on Tahiti or Moorea. The result is that the influence of runoff on lagoon sedimentation be­­ comes almost negligible, practically all the sedimentary particles being calcareous, derived from marine organisms (Guilcher et al. 1969). There is only one passage, Teavanui, more than 20 m deep, situated on the leeward side of the barrier, but the lagoon receives oceanic water by overwash of the trade wind swell in the north and east between the low islands on the barrier, and in the south when storms over the Southern Ocean generate a heavy swell that travels as far as the tropics. Overwash has built aprons of sediment, gradually filling the lagoon, but it remains fairly deep, in places more than 30 m, and the absence of rock pinnacles led to its use by the United States Navy during World War II.

7.  Maupiti The last volcano of the group that has not yet been ­completely engulfed by subsidence resembles a smaller ­Bora-Bora. One difference, difficult to explain, is that the lagoon is full of pinnacles. A large apron of sediment carried in by overwash is found in the north.

References Davis WM (1928) Society Islands. In the coral reef problem, Vol 9. American Geographical Society Special Publication, New York, pp 254–270 and 283–307 Guilcher A, Berthois L, Doumenge F, Michel A, Saint-Requier A, Arnold R (1969) The reefs and coral lagoons of Mopelia and Bora-Bora (Society Islands) compared with some other reefs and lagoons (Tahiti, Scilly, Western Tuamotus). (In French with English summary). ORSTOM Memoir 38, Paris

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23.4 Other Pacific Islands

Joanna Ellison

1.  Introduction Some of the Pacific larger islands (such as New Zealand, Papua New Guinea, and Japan) are described and discussed in other chapters, which also include: > Hawaii > New > Fiji

Caledonia and the Loyalty Islands

> The Society Islands, including Tahiti. For Lord Howe Island and Norfolk Island, see  > New South Wales For the Chatham Islands, see  > New Zealand For the Galapagos Islands, see  > Ecuador.

2.  Geographical Setting The divergent plate boundary of the East Pacific Rise lies to the east of the region and is the formative source of the Pacific tectonic plate. In moving to the WNW, the plate passes over a number of hot spots causing the eastern Pacific volcanic island chains of Hawaii, Pitcairn, French Polynesia, and the Cook Islands, and island chains further to the west such as the Samoas are also volcanic hotspot remnants. The Pacific plate meets and mostly subducts beneath the Australian plate to the east of the main islands of Tonga, Fiji, and Vanuatu and north of the Solomons and Papua New Guinea. To the east, all islands are arisen through the ocean, whereas to the west some larger islands are migrated Gondwanan fragments. This has a great influence on species biodiversity, with true islands being low in species with those present are adapted to long-distance oceanic migration. Island types and geological ages and presence or absence of rivers were summarized by Resh and De Szalay (1995). This reflects a progression of island types as first described by Darwin (1840), volcanic islands later develop fringing coral reef margins, and start to erode and submerge. As this progresses, coastal plains and ­mangroves and

freshwater swamps are likely to develop. Fringing reefs become barrier reefs with further island submergence, increasing habitat variety in reef environments and introducing lagoons and seagrass habitats. Further subsidence leads to atoll development with loss of the central volcano. This progression was illustrated by Chubb (1957) for a number of island chains in which “vulcanicity began at one end and moved progressively along the chain, erecting a series of volcanoes, each of which passed through a succession of stages, ending as an atoll.” He cited a number of atolls in the NW Society Islands. Southeastward from these, there are Maupiti, a near-atoll, then Bora-Bora, a deeply dissected and embayed volcano with a barrier reef and broad lagoon; then progressively the less dissected and less embayed volcanoes of Tahaa, Raiatea, Huahine, Moorea, and Tahiti, the last consisting of two volcanoes, both still showing their original volcanic form and having an almost unindented coastline. All these islands have barrier reefs, which on Tahaa extend nearly two miles offshore, on the succeeding islands generally to about 2 km, and around parts of Tahiti to only 1 km, becoming a fringing reef in places. Finally, at the eastern end of the chain, lies Mehetia, a bold volcanic cone, hardly dissected and unembayed, with a narrow fringing-reef and a well-­ preserved crater from which many recent lava-flows have issued, though there is no record of volcanic activity within human memory. Other examples given by Chubb include the Hawaiian chain, which becomes older northwestward, the Samoan chain, which becomes older southeastward, and the older Tuamotu (Mangareva to younger Pitcairn) and eastern Caroline chains (Chuuk, Ponape to the youngest Kosrae).

3.  Volcanic Islands Most of the active volcanoes are located on the active volcanic arcs along the western margin of the Pacific Ocean, such as the island arcs of the Solomon Islands and Vanuatu. To the east of the Pacific, active volcanoes occur on three

A revised version of a chapter by J.C. Schofield in The World’s Coastline (1985: 1027–1033). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_23.4, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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isolated groups: the islands of Hawaii and Maui at the southeast end of the Hawaiian chain of islands (>Hawaii), Revilla Gigedo west of Mexico, the Galapagos Islands near the Equator on the eastern margin of the Pacific Basin, where Darwin considered there could be 2,000 craters. Tonga is located on the Australian plate immediately west of the Pacific plate subduction trench. About 200 km west of the subduction trench active volcanoes occur, including the larger volcanic islands of Niuatoputapu, Niuafo’ou, Fonualei, Late, Kao, Tofua, and Fonuafo’ou (Park and Whistler 2001). The most youthful or shoal stage of active volcanos are typified by Hunga Ha’apei about 50 km NW of the main island of Tongatapu, from which rafts of pumice float toward the beaches of adjacent islands, and indeed as far west as north Queensland. Other volcanoes, probably dormant, occur on Uvea in the Wallis Islands northeast of Fiji, and Mehetia at the southeast end of French Polynesia. Volcanic coasts usually reflect long histories of development and are steep and rocky. Some of the older coasts, particularly those exposed to the prevailing winds, have been cut back to form high cliffs (megacliffs), commonly more than 300 m high, for example, the active volcano of Raoul Island in the Kermadec Group. Where coastal retreat has been faster than stream downcutting, the cliffs feature hanging valleys and waterfalls. Some steep, straight, or curvilinear coasts may be due to faulting or slumping, as along the southeast coast of Upolu in Samoa (Kear and Wood 1959). Where relatively young lava flows project beyond the older parts of the coast, they form protective headlands. The newness of the coast around an active volcano and the easily eroded sediment inhibit coral reef growth. How­ ever, coral, foraminiferal, and molluscan fragments are important constituents in some beaches, such as those on the leeward side of Hawaii. Other beaches, where coastal erosion is active or where streams export large quantities of sediment, are dark with volcanic fragments. The steep offshore slope carries eroded material out of the coastal system, and thus prograded coasts are absent. Isla San Benedicto is one of four widely spaced islands in the Islas Revilla Gugedo west of Manzanillo, Mexico. It is of recent volcanic origin, with a cone of pyroclastic debris dissected by vertical runoff rills and undercut by marine cliffs on the exposed western coast. Easter Island (Rapa Nui) is a roughly triangular volcanic island with an area of 120 km2, lying 3,600 km west of the Chilean coastline. The climate is subtropical, with a rainy winter and a dry summer: easterly winds prevail in spring and summer, westerlies in winter. The island has several extinct volcanoes, the largest of which is Terevaka

(601 m) in the NW, while Rano Kau (410 m) in the SW has only a narrow crater rim on the southern side, with cliffs of lava plunging into deep water. Much of the coastline is steep and cliffed, with irregular black rocky promontories and a few outlying islets off the south-west coast, such as Moto Kankau, Moto Nui, and the tall pinnacle of Moto Iti, all guano-coated bird refuges. There are no shore platforms or coral reefs, and the cliffs plunge far below sea level. Basaltic lava predominates, but there are outcrops of  red slag and the volcanic ash from which the famous ­monuments were carved. There are many incised ravines, mostly dry, some descending to rocky inlets. There are two beaches of white calcareous sand, backed by low sparsely vegetated dunes, in inlets on the north coast at Anakena and Ovame.

4.  Coral-Fringed Volcanic Islands Coral reefs are extensive in the Pacific Island region, occurring offshore of nearly all coastlines that lack significant river discharge. The absence of coral reefs around the Marquesan Islands, despite their low latitudes, could be due to an occasional influx of the cold Humboldt Current, but may reflect the eastward dwindling in the number of coral species. Nevertheless, the near-atoll of Clipperton lies still farther east and is much more isolated than the Marquesas but not reached by the Humboldt Current. The large Samoan Islands are a basaltic volcanic chain rising from an Early Cretaceous ridge 4000 m deep. The subparallel trend of the island chain with other hotspot chains on the Pacific plate suggests a hotspot origin; however, volcanism on the islands does not decline in age toward the SE. Potassium argon dates show that while Manau has the youngest lavas (Fig.  23.4.1), which is the eastern part of the Samoa Island group. The islands are of volcanic hotspot ­origin, and potassium-argon dates of the basalts of Tutuila showed active volcanism in the Early Pleistocene (McDougall, 1985). Tutuila is 32 km long, 2–9 km wide, with steep and rugged terrain descending from a largely continuous central ridge of over 300 m, with maximum elevation of 653 m. Coastal topography has deep embayments and poor development of coastal plains and fringing reefs. A drowned barrier reef is found 3 km off the entrance to Pago Pago

Other Pacific Islands

  23. 4

⊡⊡ Fig. 23.4.1 Steep catchments and narrow coastal plains of Tutuila, American Samoa.

harbour, the Taema Bank, which is 10 m below present sea level at its shallowest. Living coral reefs that encircle volcanic islands have two forms: a barrier reef separated from the island by a lagoon and a fringing reef attached to the island except at the mouths of rivers. A fringing reef forms a seaward extension of a low-tide shore platform cut in volcanic formations, whereas an island with a barrier reef has an outer reef coast exposed to the full force of the ocean, particularly active on the windward side; an inner, lagoonal reef coast, and an innermost, lagoonal volcanic coastline. Both barrier and fringing reefs provide a living, renewable defence against coastal erosion, softening the effects of tsunamis and stormy seas generated by hurricanes. Coral reefs also enable more sheltered ­communities of seagrass and mangrove to develop in their lee. Fringing coral reefs and barrier lagoons, particularly the calcareous algae that grows in association with reefs, are an important source of calcareous beach sand. For example, not only are the beaches around the high volcanic island of Rarotonga formed of marine carbonate sands, but these continue inland as a continuous vegetated strip of carbonate sand up to a height of 8 m and commonly 200–500 m wide (Wood and Hay 1970). Similar relationships are shown on the geological maps for Samoa (Kear and Wood 1959), while on Lord Howe

Island in the Tasman Sea, the world’s southernmost coral reefs are a source of calcareous beach sands and Aeolian calcarenites.

5.  Limestone Islands Limestone islands mainly occur as a result of uplifted reefderived limestone, sometimes in association with volcanism such as Guam in the southern Marianas. Nauru is an elevated platform of limestone to c. 40 m, originally covered by calcium phosphate, which has now been mostly mined, reefs are fringing (> Fig.  23.4.2). Niue is raised limestone up to 60 m in height, with the topography of a raised atoll. Like the makatea limestone of the Cook Islands, the limestone is pitted and sharp. Makatea limestone is derived from a fringing reef, which is uplifted, to form a limestone wall around the volcanic core, as demonstrated by Mangaia (Ellison 1994). Mangaia is the second largest and most southerly of the Cook Islands (21°54'Σ, 157°58'Ω), with a land area of 52 km2. The island is divided into two concentric geological zones. The inner zone is a subdued basaltic volcanic cone rising to 168 m, flattened at the summit possibly by marine erosion prior to its uplift, dating between 17 and 19 million years bp. The outer zone is a complete raised limestone rim or makatea, 0.7–2 km wide, up to 70 m in

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⊡⊡ Fig. 23.4.2 Fringing reef of Nauru, with raised coral indicative of a higher Holocene sea level.

height, with erosional topography of steep terraces on the outer edge, and cliffs on the inner edge. Emerged reefs usually present a steep coastline, in some places with plunging cliffs and in others bordered by  a living reef, but everywhere rugged with karrenfeld (karstic) solution features. A wave-cut or solution notch is common at the base of the cliff and on some cliffs is preserved at levels of higher oceanic stillstands. Solution has lowered the 23 m and 70 m strandlines at Niue by 3 m and perhaps 15 m, respectively (Schofield 1959), but is unlikely to have constructed the interior basin of this emerged atoll. This basin contains uncemented carbonate sand with Plio-Pleistocene lagoonal faunas that probably lived well below tide mark. The theory that such basins have been formed solely through solution (Purdy 1974) depends on the comparison with solution effects on blocks of non-pervious limestone. In fact, emerged reefs are extremely permeable, and there is no differential runoff of regional significance necessary for such a theory. The rain soaks straight into the ground, and any lowering of the surface is likely to be fairly general, with some more cemented parts such as beach rock and cemented boulder ramparts possibly being preferentially preserved as relict or shadow structures (Schofield 1959). However, solution at the contact between a fringing reef and less impermeable weathered volcanic rock, from which there is runoff, results in the devel­ opment of swampy hollows that are much favored for

g­ rowing taro. Such might even be the origin emerged makatea barrier reef at Mangaia, 70 m high, which may, before taro-swamp solution, have been a fringing reef (Wood and Hay 1970). In Tonga, the larger islands of ‘Eua, Tongatapu, Ha’apai, and Vava’u are formed of uplifted limestone, lying on the edge of the Australian plate to the east of the subduction trench. Tongatapu is composed of raised limestone of Quaternary origins, with Early Pleistocene uplift forming the 25–65 m south coast ridge, and subsequent leeward sedimentary infill linking former patch reefs to enclose the shallow central Fanga ‘Uta lagoon. Volcanic ash has enriched the soils of much of the island. While coastal accretion continued to the northwest, erosion occurred on the southeast coastline, causing the present day rugged cliff topography. A 7 m narrow terrace in these cliffs formed during the last interglacial. Evidences of a higher Holocene sea level stand occur below this terrace, such as notches, algal platforms, and sea caves at around 2 m above present sea level. Isotopic dates of this mid-Holocene highstand have been carried out on associated features on the north shore of Tongatapu, at which time the western arm of the Fanga ‘Uta lagoon could freely exchange with the ocean through a channel where the capital, Nuku’alofa, is now located. Coral in-growth position at this site gave uranium–thorium and radiocarbon dates indicating a former sea level 2.2 ± 0.3 m.

Other Pacific Islands

To the south of Tongatapu, ‘Eua is derived from Gond­ wanan volcanic fragments in association with uplifted limestone, which are dissected by caves. The island rises to  312 m, and spectacular cliffs occur on the windward western shore.

6.  Atolls Some countries in Micronesia and Polynesia are almost entirely composed of coral atolls, including Kiribati, Tuvalu, and the Marshall Islands. There are 13,000 km2 coral reefs in the countries east of Samoa when compared with 6,000 km2 of land. The Republic of the Marshall Islands consists of 29  coral atolls and five non-atoll coral islands spread over an exclusive economic zone of over 2 million km2 of ocean between 4 and 15° N. With just 181 km2 of land, all 607 islands are low lying, mostly less that 2 m above MSL. The islands form two almost parallel chains, Ratak (sunrise) to the east and Ralik (sunset) to the west. Many islands are very small but the Marshalls atolls typically are very large in lagoonal area, Kwajalien has the largest atoll lagoon in the world. Ailuk (10° 58' N, 169° 88' E) is one of the more remote of the 29 atolls, with a continuous reef surrounding a

⊡⊡ Fig. 23.4.3 Windward reef flat of Ailuk Atoll, Marshall Islands, showing aragonite cementation of reef flat and storm beach deposits.

  23. 4

lagoon of 177 km2, and 55 islands of only 5.36 km2 land inhabited by about 400 people. The lagoon is deep and has a continuous encircling reef with four through channels. The lagoon water is quite deep with rich coral biodiversity. An atoll, that ‘most extraordinary of natural structures’ (Davis 1928), consists of a coral reef surrounding and protecting a usually tranquil lagoon, and supporting small vegetated islets 3–4 m high. The opalescent light greens and blues of the lagoon are in marked contrast to the startling white beaches, deep palm-green islands, algal-red outer reef, and deep clear blue ocean. Chemical solution and redeposition of calcite and aragonite cement are an important agency in beach protection. Cementing of beach sand to form beach rock and of storm-built boulder banks to form boulder ramparts is particularly active within the intertidal zone, and can be very rapid (>Fig. 23.4.3). Isolated reef blocks can project above the general level of the reef platform. They may represent remnants of a higher reef, particularly if present on the leeward side of an atoll, or they may be large boulders torn from the reef front during a cyclone, which can be distinguished by not being in growth position (> Fig.  23.4.4). In the latter case, they are perhaps the remains of a much more extensive storm ridge, the smaller blocks having been removed before the ridge could be cemented to the reef flat.

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⊡⊡ Fig. 23.4.4 Storm-deposited coral boulder on Ailuk Atoll.

7.  Islands on Coral Reefs Islets (sand cays, ridges of coralline gravel) on coral reefs owe their preservation from all but the greatest storms to the presence of cemented boulder ramparts on their windward sides, partly aided by beach rock. In some islets, multiple parallel ramparts may form much of their width. In others, their greater width consists of unconsolidated sand on the leeward side of the rampart. The predominance of islets on the windward reefs is probably due to a more common construction of boulder ramparts there. Mangroves may fringe the leeward shore of lagoons. Skeletal soils and radiocarbon dates (Schofield 1977) suggest that the islets are late Holocene in age, and the relationship between Late Holocene sea level fall and islet formation is debatable. There is little evidence in the Marshall Islands for a Holocene sea level much higher than present while Schofield (1980) demonstrated a maximum of + 2.4 m for Kiribati almost 3,000 years bp. Quantitative and mineralogical studies of prograded coasts in New Zealand show that 95% or more of the coastal sand has been derived from the sea floor as a result of the 2 m fall in sea level in the late Holocene, and this is also probably true for Pacific reef islets. Even those on the windward reef are built from distinctive lagoonal sand (Schofield 1977), but this is to be expected, as the wide, shallow, marginal lagoonal floors are the only source from which sand could be derived as a result of sea level fall.

The preservation of more or less permanent islets has almost certainly been favored by the net fall in sea level since the maximum of the Late Quaternary marine transgression. This sea level fall took place about 3,000 years bp, and it may be no coincidence that this is also about the earliest known Polynesian settlement date within the Pacific Ocean. Reef islands and atolls include broad low-tide shore platforms. The outermost zone is often a 0.5–1 m high algal ridge, alive in constant spray at the breaker line. Deep grooves running across it give a saw-toothed appearance to the edge of the reef. Inside the ridge in some places, there is the so-called boat channel or string of deep pools that persist even at low tide, interconnected with the ocean, and support live corals. Landward, the platform proper may consist of an older reef and its outer algal ridge, which have been abraded flat, but in which the saw-toothed nature can still be recognized from the air: presumably as the algal ridge was abraded the cross grooves have been infilled.

References Chubb LJ (1957) The pattern of some Pacific island chains. Geol Mag 94:221–228 Darwin C (1840) Journal of researches into the geology and natural history of the various countries visited by H.M.S. Beagle from 1832 to 1836. H. Colburn, London

Other Pacific Islands Davis WM (1928) The coral reef problem. American Geographical Society, New York (Special Publication 9) Ellison JC (1994) Caves and speleogenesis of Mangaia, Cook Islands. Atoll Res Bull 417:1–25 Kear D, Wood BL (1959) The geology and hydrology of Western Samoa. N Z Geol Surv Bull 63 McDougall I (1985) Age and evolution of the volcanoes of Tutuila, American Samoa. Pac Sci 39:311–320 Park G, Whistler A (2001) The terrestrial ecology and botany of Tofua and Kao Islands in Ha’apai, Kingdom of Tonga. South Pacific Biodiversity Conservation Programme, Apia, p 136 Purdy EG (1974) Reef configurations, cause and effect. In: Laporte LE (ed) Reefs in time and space. Society of economic paleontologists and mineralogists, Special Publication 18, pp 9–76

  23. 4

Resh VH, De Szalay FA (1995) Streams and rivers of Oceania. In: Cushing CE, Cummins KW, Minshall GW (eds). River and stream ecosystems. Elsevier Science, Amsterdam, the Netherlands, pp 717–739 Schofield JC (1959) The geology and hydrology of Niue Island, South Pacific. N Z Geol Surv Bull 62 Schofield JC (1977) Effect of late holocene sea-level fall on atoll development. N Z J Geol Geophys 20(3):531–536 Schofield JC (1980) Postglacial transgressive maxima and second-order transgressions of the south-west Pacific Ocean. In: Morner N (ed) Earth rheology. Isostasy and Eustasy. John Wiley & Sons, Chichester, pp 517–521 Wood BL, Hay RF (1970) Geology of the cook islands. N Z Geol Surv Bull 82

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24.1 Antarctic Coast 

1.  Introduction The Antarctic coastline is just over 30,000 km long. Apart from studies of the ice-fringed coasts there have been few geomorphological studies of Antarctic coastal landforms, and work has understandably clustered around the locations of scientific bases in the Antarctic Peninsula and the Ross Sea. There has also been a preoccupation with the relationship between ancient coastlines and the glacial history of the continent. Geomorphological information on the relatively limited sectors of rocky and beachfringed coastline is still very sparse. The coastline can be examined on Google Earth, and details of coastal segments can be seen on http://visibleearth.gov/view, but it is difficult to provide an along-the-coast account of the kind presented in other chapters. The present chapter is based on a review by Kirk (1985), advice from Janet Thomson and Adrian J. Fox of the British Antarctic Survey, and information obtained from Internet sources and from Stonehouse (2002). The Antarctic region comprises the great southern continent, its surrounding cold, ice-strewn ocean, and associated islands. It is customarily thought of as lying south of latitude 66° 30ʹ (the Antarctic Circle), but this line intersects the continent in sectors between 50° and 60° E (Napier Mountains) and 90° and 140° E (Wilkes Land). The Antarctic continent has an area of 13.1 million sq. km. It has a rounded outline, with a long, narrow winding mountainous spur, the Antarctic Peninsula, extending towards > Argentina. Massive ice sheets cover most of the continent, and are bordered by ice shelves of varying width, the largest being the Ross Ice Shelf and the Ronne Ice Shelf in broad embayments facing the Weddell Sea in the northwest and the Ross Sea on the southern side. A line joining these two embayments divides West Antarctica from East Antarctica, and is bordered by the Transantarctic Mountains. East Antarctica occupies ­three-quarters of the area of the continent and comprises a vast, domed ice sheet rising to an elevation of more

than 4,000 m and having only small ice-free areas at the coast, many of which occur along the western Ross Sea where the Victoria Land mountains form the shore. By contrast, the much smaller area of West Antarctica can be thought of as a large archipelago submerged in ice. Much of its ice-free coastal area is mountainous and occurs in the Antarctic Peninsula. The Antarctic Treaty (1961) established international co-operation and freedom of scientific investigation in the continent. Several sectors of the Antarctic coastline have been claimed (some overlapping) as follows (anticlockwise sequence from 150° W): The Ross Dependency (New Zealand), Australian Antarctic Territory, Adélie Land (France), Australian Antarctic Territory, Queen Maud Land (Norway), British Antarctic Territory, Chilean Antarctic Territory and Argentine Antarctica. A major unclaimed territory lies between 90° and 150° W. The Antarctic is snowbound in winter. Only about 0.33% of the continent has land exposed during the summer (in the Antarctic Peninsula this increases to 3.93%). Exposed coastal rock areas are subject to periglacial (freeze-and-thaw) shattering, and unconsolidated surface deposits are modified by freeze and thaw, which develops patterned ground (stone stripes, polygons) on areas that remain undisturbed, with permafrost at depth. In some ice-free areas there has been accumulation of peat-moss deposits up to 1.5 m thick over the past 7,000 years, again with permafrost at depth. The Antarctic coastline has the following divisions: 44% is the margin of shelf ice, 38% ice walls, 13% glacier snouts and only 5% rocky or beach-fringed (Drewry et al. 1982) (>Fig.  24.1.1). Most of the coast thus consists of ice  cliffs at the margins of either floating (shelf) ice or grounded ice sheets. The coastline represents the contact between oceanic water and continental ice, and generally has a smoothly curved outline, although it is known from ­seismic and other survey work that fiords and rocky archipelagoes occur beneath the ice. Shelf ice calves into icebergs that drift away northward and gradually melt, while the surrounding waters have sea

A revised version of a chapter by R.M. Kirk in The World’s Coastline (1985: 1047–1053). Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_24.1, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Antarctic Coast

⊡⊡ Fig. 24.1.1 The coastline of Oates Land, Australian Antarctic Territory, showing a glacier reaching the shore on a steep rocky sector, with an outlying iceberg and ice floes. (Courtesy Australian Antarctic Division, photograph by P.G. Law.)

ice which increases in area from 4 million sq. km in late ­summer (February) to 20 million sq. km in late winter (September). The Antarctic climate is cold and dry, and katabatic winds blow off the continent at velocities of up to 300 km/h. Inshore waters are either calm (protected by sea ice) or have only occasional incoming waves to shape beaches. It is thought that the Antarctic ice shelves have formed largely by the seaward movement of ice that formed over the land, rather than by continued thickening of sea ice bordering the shore. Like the grounded ice sheets and glaciers they contain sediment, and the upper surfaces of icebergs derived from them become coated with sediment as they melt. Much attention has been given to the nature and dynamics of the ice coasts (Williams et al. 1998). The coastline of floating ice margins can advance and retreat by several kilometres in a decade, whereas the margins of grounded ice sheets are relatively stable. Ice shelves hold back the outflow of glaciers, which will accelerate if they disappear. The United States Geological Survey has been mapping changes in the area and volume of Antarctic ice sheets (Swithinbank et al. 1996). Disintegration of the Wordie Ice Shelf on the Antarctic Peninsula between 1936 and 1992 was documented by the British Antarctic Survey (1996) and in 2000 this agency, in association with the United States Geological Survey, began mapping coastline changes and glacial features on

the Antarctic Peninsula. Further losses have occurred, including the separation of a 570 sq. km section of the Wilkins ice shelf in March 2008. There is much concern about the effects of Antarctic ice reduction on global sea level, for if all the grounded ice in Antarctica were to melt there would be a rise of about 70 m in the oceans (the melting of sea ice and ice shelves has little effect on ocean volume). Further information can be obtained from the United States Geological Survey on http://chht-ntsrv.er .usgs.gov/projects/ere.glaciers.html and http://chht-ntsrv .er.usgs.gov/projects/ere.glaciers.html, and from the British Antarctic Survey on http://www.antarctica.ac.uk. During the summer the reduction of sea ice exposes a few sectors of rocky and beach-fringed Antarctic coastline, to incoming wave action and marine processes, particularly on the Antarctic Peninsula and outlying islands, where there are steep and cliffy coasts, broken by fiords, dissected into caves, stacks and islands, and in places fronted by shore platforms. Generally the ice cliffs descend to an exposed shore of glaciated rock outcrops, or to morainic fringes partly reworked to beaches by summer wave action. Sometimes the basal rock outcrop is a horizontal bench which may be a strandflat of the kind that formed beneath ice sheets on the coast of Scandinavia (> Norway), developing as the result of recurrent freezeand-thaw processes beneath the ice. Further seismic profiling of the subglacial topography is needed to determine whether there is a strandflat here.

Antarctic Coast

Beaches are derived mainly from glacial drift deposited on or near the coast, and to a lesser extent to erosion of rocky coast outcrops. On the Antarctic Peninsula ground moraine is thought to be the most important source of beach materials, but on the Ross Sea coast the products of cliff abrasion, including talus and wind-blown sediment, are the predominant sources of beach material, though ground moraine is locally important at higher latitudes. Antarctic beaches differ from those of temperate latitudes chiefly in the seasonal effects of ice. There is little difference (apart from the relative levels of the land and sea and the extent of ice cover) between existing conditions and those that obtained throughout the Quaternary. The ice sheets erode and transport sediment of a wide variety of sizes and shapes and where these are delivered to the shore they are subject to the highly selective action of waves during the summer thaw, these processes being excluded by the formation of sea ice in winter. Antarctic beaches are distinctive in that they rest upon ice or permafrost, are pitted where ice blocks have melted, and show ridges and mounds formed by ice-push and/or deposition from stranded sea ice. Ridges formed by ice-push from wind-drifted or current-drifted sea ice commonly show beach scars. Wave-built berms may end laterally where ice was present at the time of their formation, or where there are erosional gaps formed by meltwater outflow. Ice-rafted rock fragments and glacial erratics are often found on beaches, with striations formed by grounding sea ice and icebergs. Most beaches are associated with ice-contact features (proglacial deltas, proglacial ramparts, and esker-like features) and glaciomarine sediment, and lobes or aprons of solifluction debris spread on to them during summer. Pebbles are usually poorly rounded, and ventifracts (wind-scoured pebbles) are common. Frost cracks and mounds, stone circles and polygons form as the result of freeze-and-thaw processes, and persist until the beach is disturbed by wave action. Beaches may contain cold-water fossils and the soft parts of marine organisms, as well as bones, notably whale bones. Longshore drifting occurs when waves arrive at an angle to the coast during the brief summer thaw. This can result in the formation of spits, forelands and prograded beaches, with successions of wave-piled beach ridges in patterns that indicate stages in growth. Because Antarctic beaches are commonly the nesting sites of large numbers of penguins, beach ridges are often cemented by appreciable thicknesses of guano, and the swales are devoid of pebbles because they have been gathered for nesting materials.

24.1

Extensive beach ridges occur at several sites around the Ross Sea, in the Antarctic Peninsula and on several of the offshore islands. Depositional forelands have ­commonly accumulated where longshore drifting has delivered beach material to sites that lie to leeward of steep, eroding peninsulas or the cliffed coasts of islands. Emerged (raised) beaches and beach ridges are common, associated with emerged shore platforms. These have received much attention because of their perceived significance in unravelling the Quaternary glacial history of the continent and its relationship with changing world climates and sea levels. Catkin and Nichols (1972) noted that emerged beaches have been found all round the Antarctic continent, and that some predate the most recent glacial episode, although most are younger. They commonly occur at elevations of less than 60 m above sea level, though higher areas exist. On the Soya coast in East Antarctica there is evidence up to 16 m of emergence over the last 6,000 years, with beaches under 6 m in elevation thought to be younger than 500 years (Omoto 1977). On the other hand, emerged beaches 6 m above sea level bearing boulder pavements in the Vestfold Hills are thought to be about 6,000 years old (Peterson et al. 1988). The western Ross Sea coast has emerged beach ridges dating to 5,000 years bp occur at elevations up to 20 m above sea level. Nichols (1969) found more than 20 well-developed beach ridges occur up to 34 m above sea level at Marguerite Bay on the southwest coast of the Antarctic Peninsula and concluded that the oldest known raised ridges in Antarctica are probably of Late Pliocene Age, whereas the youngest are Late Holocene. Little is known of the spatial patterns of emerged beach heights even in the McMurdo Sound region of the Ross Sea, where the glacial history of the later Quaternary is now well known. Emerged beaches therefore present interesting problems concerning glacioisostacy and glacio-eustacy and their existence in large numbers at appreciable altitudes on most sectors of the Antarctic coast argues that in the recent past the rate of isostatic uplift has exceeded the rate of sea level rise that resulted from retreat of the ice sheets. Because of the past and present extent of the ice cover the continental crust of the Antarctic is considerably depressed. Most of the coast has been extensively glaciated. The Antarctic continental shelf has an area of about 4 million sq. km, some 60% (2.4 million sq. km) of which is ice-free. The shelf has a mean depth of 350 m (almost twice that of other continental shelves), a mean width of 128 km, and a minimum width of 48 km. The continental shelf is widest beneath the Weddell Sea (611 km) and the Ross Sea (1,127 km). These characteristics, together with

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Antarctic Coast

the problems posed by stormy seas, drifting pack ice and icebergs, pose major constraints on the search for exploitable Antarctic resources, particularly offshore oil and gas reserves.

2.  The Antarctic Coastline The Greenwich meridian (0°) crosses the Antarctic coastline in Dronning Maud Land, and at this point the ice shelf is about 200 km wide. To the west spurs of the Kronprinsesse Martha coast project across it, but then it widens to the Brunt Ice Shelf off the Caird coast. Grounded ice extends to the east coast of the Weddell Sea, the southern part of which is occupied by the broad Filchner and Ronne ice shelves on either side of Berkner Island. To the west is the mountainous Antarctic Peninsula, where a narrow ice shelf extends out from the numerous fiords of Palmer Land, then broadens as the Larsen Ice Shelf. To the north the Nordenskjold coast is free of ice in summer, and has steep mountain slopes descending to the sea, as on James Ross Island. To the northeast Seymour Island has badland ravines in Tertiary sediment discharging muddy sediment from ravines in the southern cliffs (icebound in summer), and an estuarine inlet beside Cape Lamas full of mud. The high Antarctic Peninsula breaks up into islands, including Joinville Island to the north and Dundee Island, a circular volcanic cone, to the east. Lemaire Channel, passing behind Booth Island, is bordered by precipitous cliffs.

Offshore to the west are the South Shetland Islands (see below). Ice shelves reappear south of Adelaide Island, and ice  extends through George VI Sound behind Alexander Island. On the southern side is the 150 km wide Wilkins Ice Shelf, extending out to Charcot Island and Latady Island. Ice shelves occupy most of the coast of the Bellinghausen Sea and Amundsen Sea bordering West Antarctica, Siple Island being attached to the mainland by the Getz Ice Shelf. On a larger scale the Ross Ice Shelf, up to 750 km wide, occupies a bay behind Ross Sea, and the mountains of Victoria Land form an irregular eastern coast with projecting ice tongues, and are truncated by the Oates Coast. Beach ridges and a cuspate foreland occur on Cape Adare (> Fig. 24.1.2). The ice sheets of George V Land, Terre Adélie and Wilkes Land terminate seaward in ice cliffs with segments of shelf ice, and beyond the Australian base at Casey are the broader Shackleton Ice Shelf and West Ice Shelf on the shores of Davis Sea. Near the Australian base at Davis the Vestfold Hills form a peninsula, and behind Prydz Bay the Amery Ice Shelf occupies a broad re-entrant. To the north the ice cliffs are interrupted by the rocky megacliffs of Scullin Monolith, plunging into the sea from a height of 450 m. North of the Mawson base the coast curves round to Enderby Land, where there are segments of steep rocky coast, and these recur on the Prins Olav Kyst to the west. Beyond the Prins Harald peninsula are the wider Prinsesse ⊡⊡ Fig. 24.1.2 A large depositional foreland and beach-ridge complex at Cape Adare, North Victoria Land. The ridges are colonised by Adelie penguins and separated by melt pools. The complex extends for 2 km left to right and has accumulated from longshore-drifted abrasion products in the lee of the peninsula behind.

Antarctic Coast

Ragnhild and Prinsesse Astrid ice shelves, backed by glaciers fringing the rocky mountains of Dronning Maud Land and extending west past the Greenwich meridian.

3.  Subantarctic Islands Antarctica is separated from the other continents by a vast ocean belt with an area of 37.8 million sq. km and a mean depth of 4,000 m. Dotted through this ocean, and lying for the most part near the continent, are a series of mainly small islands having a total area of about 38,000 sq. km. Though separated by very large distances, these islands present distinctly cliffed coasts with small associated beaches. They have many other features in common, since most are of volcanic origin (the emergent parts of volcanic ridges formed along diverging fractures in the ocean floor crust) and their coasts rise steeply from deep water to high elevations. Many support ice caps or permanent snow fields, and their precipitous shores are commonly subjected to very heavy surf action together with the erosional effects of ice. Generally the highest sea cliffs are on the western coasts, exposed to strong waves generated by the prevailing westerly winds and storms. The larger subantarctic islands (mainly below latitude 50° S, but excluding those associated with southernmost South America) are listed in counterclockwise sequence from the Greenwich meridian. The South Sandwich Islands (56°–60° S, 28° W) (British) are a group of steep, cliffy volcanic islands with an 80% ice cover. There has been volcanic activity within the past century. South Georgia Island (53°–55° S, 39° W) (British) is a curved island 170 km long and up to 40 km wide. It has a rugged topography of glaciated mountains and retains a 57% ice cover. Deep glacially scoured valleys descend to fiords, mainly on the north coast. Shag Rocks (55° S, 42° W) (British) are a group of volcanic pinnacles up to 71 m high with precipitous plunging cliffs. The South Orkney Islands (60° S, 45° W) (British) consist of four large and several small islands. The largest, Coronation Island, is mountainous, rising to 1,265 m on Mount Nivea. The coasts are steep and cliffy, embayed, with a few beaches. The South Shetland Islands (61°–64° S, 54°–63° W) (British) are a chain of volcanic islands parallel to the west coast of the Antarctic Peninsula. They are mountainous, with an ice cover of about 80%. Clarence and Elephant Islands, to the northeast, have extensive moss peat up to 3 m thick. King George Island, the largest, has a 90% ice

24.1

cover, but its east coast, sheltered from prevailing westerly storms, is embayed, Admiralty Bay and Maxwell Bay having low-lying fringes utilised by many national Antarctic Science stations. Deception Island, to the southeast, is a partly collapsed volcanic cone, with a rim rising to 580 m overlooking a deep rounded caldera penetrated by the sea (similar to Santorini in Greece and Krakatau in Indonesia). There have been several volcanic eruptions here, notably in 1967, 1969, 1970 and 1991–1992. Earlier, in 1930, the caldera floor subsided 5 m during an earthquake. To the south, the Byers Peninsula on Livingston Island has beaches backed by broad emerged beach ridge plains (British Antarctic Survey 1996). There are also wide shore platforms cut across dipping strata by wave action and sea ice abrasion (Hansom and Kirk 1989). The Norwegian island, Peter the First Øy (69° S, 91° W), is an extinct volcano rising 1,640 m above the sea, almost completely ice-covered and pack ice fringed. Between protruding glacial lobes the coastline has a few narrow rocky bays with gravel beaches. Scott Island (67° S, 179° E) (New Zealand) is a small island, a dissected volcanic cone with an outlying stack, Haggit’s Pillar, to the west. Antipodes Islands (50° S, 179° E) (New Zealand) are a group of small cliffy islands of volcanic rock, with peat and grassland. There is an exceptionally large and varied population of nesting seabirds. Campbell Islands (52° S, 169° E) (New Zealand) are a dissected volcano, the main island rising steeply to Mount Honey (569 m). Auckland Islands (50° S, 166° E) (New Zealand) consist of a main island with fiords on its east coast and steep slope-over-wall topography on its more exposed west coast, where a periglaciated slope is undercut by marine cliffs. Balleny Islands (66° S, 163° E) (New Zealand), situated across the Antarctic Circle and 200 km north of the Oates coast, are volcanic peaks with abrupt basalt cliffs and only very small beaches. Macquarie Island (54° S, 158° E) (Australia), 1,460 km SSE of Tasmania, is 34 km from north to south and up to 5 km wide. It has steep slope-over-wall coasts rising to an undulating tussock grass peaty plateau 240–340 m above sea level, dotted with lakes. Free of snow in summer, the island has almost continuous strong westerly winds, and strong wave action has cut higher cliffs along the more exposed western coast. The east coast, almost straight, is dominated by a periglaciated slope extending down to low basal cliffs. At the northern end a low gravelly isthmus, bordered by shingle beaches, links Wireless Hill to Macquarie Island. Heard Island (53° S, 73° E) (Australia), 3,700 km southwest of Perth (Western Australia), is a volcanic island on

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Antarctic Coast

⊡⊡ Fig. 24.1.3 Subantarctic Heard Island is a volcanic cone, and recurrent vulcanicity supplies abundant sand and gravel, which is carried down to the coast in glaciers. Very strong westerly winds produce waves that sweep beach material along the northern and southern coastlines to accumulate in a large trailing spit at the eastern end of the island. (Courtesy Geostudies.)

the Kerguelen-Gaussberg Ridge. It rises to Mawson Peak (2,745 m), summit of the Big Ben volcano which erupted in 1950, 1985 and 1992. Small glaciers radiate from the cone, and deliver morainic material to the coast, forming beaches and a spit at the eastern end (> Fig. 24.1.3). McDonald Islands (53° S, 72° E) (Australia), are small volcanic islands 40 km west of Heard Island, also on the Kerguelen-Gaussberg Ridge, about 4,500 km SW of Perth. After a prolonged dormant period it has been expanding as the result of volcanic activity since 1992, and has absorbed several former islets. It has doubled in area since it was surveyed by helicopter in 1980. There is little or no vegetation, but penguins have colonised. There are beaches of black volcanic sand and pumice. Iles Kerguelen (49° S, 70° E) (France), towards the northern end of the Kerguelen-Gaussberg Ridge, consist of one large island (Grande Terre) and many small rocks and islets (Guilcher 1972). It has a 10% ice cover (Cook Glacier), and was formerly strongly glaciated, with valleys descending to deep fiords on the northern, eastern and southern coasts.

The west coast is steep, cliffed and relatively straight. Bouvet Island (Norway) (54° S, 3° E) is a small and remote volcanic island 1,600 km SSW of Cape Town, with many islets and rocks, projecting from a submarine ridge that runs from the mid South Atlantic into the southwest Indian Ocean. It rises to a peak 780 m above sea level with an ice-filled crater, and the surface is more than 90% icecovered. Glaciers descend to ice cliffs on the east and south coast, while there are basalt cliffs up to 490 m high on the exposed west coast. A lava shelf formed on the west coast in the 1950s. Thompson Island, charted in the nineteenth century to the northeast, is thought to have disappeared as the result of a volcanic explosion.

References British Antarctic Survey (1996) Geomorphological map of Byers Peninsula, Livingston Island, Cambridge Catkin PE, Nichols RL (1972) Quaternary studies in Antarctica. In Adie RJ (ed) Antarctic Geology and Geophysics Symposium, Univer­ sitetsforlaget, Oslo

Antarctic Coast Drewry DJ, Jordan SR, Jankowski E (1982) Measured properties of the Antarctic ice sheet. Ann Glaciol 3:15–26 Guilcher A (1972) Contribution a l’étude géologique et géomorpholgique des Iles Kerguelen. Mar Geol 13:289–290 Kirk RM (1985) Antarctica. In: Bird ECF, Schwartz ML (eds) The World’s Coastline. Van Nostrand Reinhold, New York, pp 1047–1053 Nichols RL (1969) Geomorphic features of Antarctica. In Bushnell VC (ed) Antarctic Map Folio Series. Folio 12, Geologic Map of Antarctica, American Geographical Society, New York Omoto K (1977) Geomorphic development of the Soya Coast, East Antarctica-chronological interpretation of raised beaches based on levellings and radiocarbon datings. Scientific Reports, Tokohu University, 7th Series. (Geography), 27(2):9 5–148

24.1

Peterson JA, Finlayson BL, Zhang Qingsong (1988) Changing dis­ tribution of late Quaternary terrestrial lacustrine and littoral ­environments in the Vestfold Hills, Antarctica. Hydrobiologia 165:221–236 Stonehouse B (ed) (2002) Encyclopedia of Antarctica and the Southern Oceans. Wiley, Chichester Swithinbank C, Williams RS, Ferrigno JG, Seekins BA, Lucchitta BK, Rosanova CE (1996) United States Geological Survey, Map I-2600-F Williams RS, Ferrigno JG, Swithinbank C, Lucchitta BK, Seekins BA, Rosanova CE (1998) Glaciological delineation of the dynamic coastline of Antarctica. Antarct J US 32:225–227

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Index Maps

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_Index Maps, © Springer Science+Business Media B.V. 2010 (Dordrecht)

3. CENTRAL AMERICA

1. U.S.A.

2. CANADA

NTIC

4. SOUTH AMERICA

5. CARIBBEAN ISLANDS

8. EUROPE

ISLA 24. ANTARCTICA

14. SOUTH AFRICA

15. EAST AFRICA

13. WEST AFRICA

18. INDIAN OCEAN ISLANDS

17. SOUTH ASIA

16. SOUTH-WEST ASIA

11. CASPIAN & ARAL SEAS

9. RUSSIAN FEDERATION

10. MIDDLE EAST

12. NORTH AFRICA

7. BRITISH ISLES

23. PACIFIC OCEAN ISLANDS

22. NEW ZEALAND

19. SOUTH-EAST ASIA

21. AUSTRALIA

20. EAST ASIA

World Map - Continents

LA 6. AT AN OCE NDS

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_Index Maps, © Springer Science+Business Media B.V. 2010 (Dordrecht)

1.16

23.3

1.1

1.3

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1.4

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3.1

2.6

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1.14 1.13 1.12 6.2 1.11 1.9 1.10 6.1 5.1 5.2 5.3 5.4 4.12 5.5 4.9 4.7

2.3 2.2 1.8

3.3 4.1

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8.2 9.1 8.4 8.5 8.6 9.6 8.29 8.28 9.5

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21.7

8.9 8.8 8.11 8.12 8.18 11.6 8.13 20.3 8.27 11.1 8.17 8.16 8.15 10.1 10.2 11.5 20.5 10.3 12.4 16.4–16.8 10.4 12.5 12.3 10.5 16.9 20.4 12.2 20.1 15.8 13.1 17.1 17.4 20.2 16.1 12.1 15.7 13.2 17.5 17.2 16.2 13.3 19.5 19.6 19.7 13.10 16.3 15.6 19.1 13.4 13.7 19.2 13.11 13.5 19.3 15.5 17.3 15.4 19.10 13.6 13.9 13.12 13.8 15.3 19.8 19.9 15.2 18.2 23.1 13.13 21.3 15.1 18.1 21.2 21.4 21.1.2 13.14 14.1 21.1 21.1.1 6.5 21.6 21.5

7.1 7.27

6.4

8.1

8.1.1

World Map

22.2

22.1

23.2

1474 Index Maps

Index Maps

Key to World Index Map USA 1.1 Alaska 1.2 Washington 1.3 Oregon 1.4 California 1.5 Texas 1.6 Louisiana 1.7 Mississippi 1.8 Alabama 1.9 Florida 1.10 Georgia 1.11 South Carolina 1.12 North Carolina 1.13 Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey) 1.14 New York and New England 1.15 Great Lakes (USA) 1.16 Hawaii

4.4 Argentina 4.5 Uruguay 4.6 Brazil 4.7 French Guiana 4.8 Surinam 4.9 Guyana 4.10 Venezuela 4.11 Caribbean Coast of Colombia 4.12 Caribbean Coasts, Panama to Belize 4.13 Mexico, Caribbean Coast

CARIBBEAN ISLANDS 5.1 Cuba 5.2 Jamaica 5.3 Hispaniola (Haiti and the Dominican Republic) 5.4 Puerto Rico 5.4.1 Virgin Islands 5.5 Lesser Antilles

CANADA ATLANTIC OCEAN ISLANDS 2.1 2.2 2.3 2.4 2.5 2.6 2.7

British Columbia New Brunswick and Nova Scotia Gulf of St. Lawrence Saint-Pierre-et-Miquelon Islands Newfoundland Northern Canada Great Lakes (Canada)

6.1 6.2 6.3 6.4 6.5

Bahamas Bermuda Greenland Iceland Other Atlantic Ocean Islands

BRITISH ISLES PACIFIC CENTRAL AMERICA 3.1 Mexico, Pacific Coast 3.2 Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama 3.3 Pacific Coast of Colombia

SOUTH AMERICA & WESTERN CARIBBEAN 4.1 Ecuador 4.2 Peru 4.3 Chile

7.2–7.23 England and Wales counties 7.24.1–7.24.10 Scottish sectors 7.25 Northern Ireland 7.26 Ireland 7.27 Channel Islands

EUROPE 8.1 Norway 8.1.1 Svalbard and Jan Mayen 8.2 Sweden 8.3 Finland

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Index Maps

1476

8.3.1 Russian Gulf of Finland 8.4 Estonia 8.5 Latvia 8.6 Lithuania 8.7 Kaliningrad 8.8 Poland 8.9 Germany 8.10 Denmark 8.11 The Netherlands 8.12 Belgium 8.13 France 8.14 Spain 8.14.1 Gibraltar 8.15 Portugal 8.16 Italy 8.16.1 Sardinia 8.16.2 Sicily 8.17 Malta 8.18 Slovenia 8.19 Croatia 8.20 Bosnia-Herzegovina 8.21 Montenegro 8.22 Albania 8.23 Greece 8.24 Crete 8.25 Bulgaria 8.26 Romania 8.27 Ukraine 8.28 Sea of Azov 8.29 Republic of Georgia

RUSSIAN FEDERATION 9.1 9.2 9.3 9.4 9.5 9.6

Russian Gulf of Finland Russian Black Sea Coast The Pacific Coast of Russia The Arctic Coast of Russia Russian Caspian Coast Kaliningrad

CASPIAN AND ARAL SEAS 11.1 11.2 11.3 11.4 11.5 11.6

Iran – Caspian Sea Coast Turkmenistan Kazakhstan Russian Caspian Coast Azerbaijan Aral Sea

NORTH AFRICA 12.1 12.2 12.3 12.4 12.5

Egypt (Mediterranean) Libya Tunisia Algeria Morocco

WEST AFRICA 13.1 Mauritania 13.2 Senegal and Gambia 13.3 Guinea Bissau 13.4 Republic of Guinea 13.5 Sierra Leone 13.6 Liberia 13.7 Ivory Coast 13.8 Ghana 13.9 Togo and Benin 13.10 Nigeria 13.11 Cameroon and Equatorial Guinea 13.12 Gabon, Congo, Cabinda and Zaïre 13.13 Angola 13.14 Namibia

SOUTH AFRICA 14.1 South Africa

MIDDLE EAST

EAST AFRICA

10.1 10.2 10.3 10.4 10.5

15.1 15.2 15.3 15.4 15.5

Turkey Cyprus Syria Lebanon Israel (with the Gaza Strip)

Mozambique Tanzania Kenya Somalia and Djibouti Eritrea

Index Maps

15.6 Sudan 15.7 Egypt, Red Sea Coast 15.8 Jordan

19.8 Indonesia 19.9 East Timor 19.10 Papua New Guinea

SOUTH-WEST ASIA

EAST ASIA

16.1 16.2 16.3 16.4 16.5 16.6 16.7 16.8 16.9

20.1 20.2 20.3 20.4 20.5

Saudi Arabia, Red Sea Coast Yemen Southern Arabia and Oman United Arab Emirates Qatar Bahrein Saudi Arabia, Persian Gulf Coast Kuwait and Iraq Iran

SOUTH ASIA 17.1 17.2 17.3 17.4 17.5

Pakistan India Sri Lanka Bangladesh Burma (Myanmar)

China Taiwan North Korea South Korea Japan

AUSTRALIA 21.1 New South Wales 21.1.1 Lord Howe Island – (New South Wales) 21.1.2 Norfolk Island 21.2 Queensland 21.3 Northern Territory 21.4 Western Australia 21.5 South Australia 21.6 Victoria Introduction 21.7 Tasmania

INDIAN OCEAN ISLANDS

NEW ZEALAND

18.1 Madagascar 18.2 Indian Ocean Islands

22.1 New Zealand 22.2 The Chatham Islands

SOUTH-EAST ASIA

PACIFIC OCEAN ISLANDS

19.1 Thailand Andaman Sea Coast 19.2 Malaysia – Introduction 19.2.1–19.2.13 Malaysian States 19.3 Singapore 19.3.1 Thailand: Gulf of Thailand Coast 19.4 Brunei (Negara Brunei Darussalam) 19.5 Cambodia 19.6 Vietnam 19.7 Philippines

23.1 23.2 23.3 23.4

New Caledonia and the Loyalty Islands Fiji Society Islands Other Pacific Islands

ANTARCTICA 24.1 Antarctic Coast

1477

1478

Index Maps

British Isles

7.24.5

7.24.6

7.24.4

7.24.8

7.24.3 7.24.7 7.24.2 7.24.1

7.24.9

7.23 7.24.10 7.22

7.25 7.2

7.21

7.4 7.3 7.5

7.20

7.26 7.6

7.18 7.7 7.8

7.9.1

7.10

7.9.2

7.19

7.17

7.7.1 7.12

7.11

7.13 7.14 7.27

7.15

7.16

Index Maps

Key to British Isles Map 7.2 Cumbria 7.3 Lancashire, Merseyside and Cheshire 7.4 The Isle of Man 7.5 North Wales and Anglesey 7.6 West Wales 7.7 South Wales 7.7.1 The Severn Estuary 7.8 Gloucestershire, Somerset and North Devon 7.9.1 The North Coast of Cornwall 7.9.2 The South Coast of Cornwall 7.10 Isles of Scilly 7.11 South Devon 7.12 Dorset 7.13 Hampshire 7.14 Isle of Wight 7.15 Sussex 7.16 Kent 7.17 Essex 7.18 Suffolk 7.19 Norfolk

7.20 Lincolnshire 7.21 Yorkshire 7.22 Durham, Tyne and Wear 7.23 Northumberland 7.24.1 South-East Scotland 7.24.2 Fife 7.24.3 Angus and Aberdeenshire 7.24.4 Fraserburgh to Moray Firth and Duncansby Head 7.24.5 The Orkney and Shetland Islands 7.24.6 North and North West Scotland (Duncansby Head to Cape Wrath and Fort William) 7.24.7 The Inner Hebrides 7.24.8 The Outer Hebrides 7.24.9 Fort William to the Clyde 7.24.10 South-West Scotland 7.25 Northern Ireland 7.26 Ireland 7.27 Channel Islands

1479

1480

Index Maps

Europe 8.1.1

9.4 6.4 8.1

8.3.1 9.1

8.3 8.2

8.4 8.5

7.24 7.25

8.6 8.10

8.7

7.26 7.2–7.23 7.27

8.11 8.12

8.9

8.8

8.13.1

8.27

8.13.2

8.13.4 8.14

8.25

8.19

8.13.3

8.15

8.26

8.18 8.20 8.21 8.22

8.16

8.16.1 8.16.2 8.17

8.23 8.24

Index Maps

Key to Europe Map 8.1 Norway 8.1.1 Svalbard and Jan Mayen 8.2 Sweden 8.3 Finland 8.3.1 Russian Gulf of Finland 8.4 Estonia 8.5 Latvia 8.6 Lithuania 8.7 Kaliningrad 8.8 Poland 8.9 Germany 8.10 Denmark 8.11 The Netherlands 8.12 Belgium 8.13.1 North Coast of France 8.13.2 West Coast of France 8.13.3 Mediterranean France 8.13.4 Corsica

8.14 Spain 8.14.1 Gibraltar 8.15 Portugal 8.16 Italy 8.16.1 Sardinia 8.16.2 Sicily 8.17 Malta 8.18 Slovenia 8.19 Croatia 8.20 Bosnia-Herzegovina 8.21 Montenegro 8.22 Albania 8.23 Greece 8.24 Crete 8.25 Bulgaria 8.26 Romania 8.27 Ukraine

1481

Index Maps

1482

W E S T

1 2

L

A

4

A

M

3

M

11

A

M S

A

I

Y

5

Y

L 10

A

E

A

S

T

12

S 9

I

6

A

7

8

Key to Malaysia Map 1 2 3 4 5 6 7

Perlis Kedah Penang Perak Selangor Negeri Sembilan Melaka

8 9 10 11 12 13

Johor Pahang Terengganu Kelantan Sarawak Sabah

A

L

A

Y

S

I

A

13

List of Entries

List of Entries Alabama 1.8 Alaska 1.1 Albania 8.22 Algeria 12.4 Angola 13.13 Angus and Aberdeenshire 7.24.3 Antarctic Coast 24.1 Aral Sea 11.6 Argentina 4.4 Atlantic Coast Central (USA) (Virginia, Maryland, Delaware and New Jersey) 1.13 Atlantic Ocean Islands – Editorial Introduction 6.0 Australia – Editorial Introduction 21.0 Azerbaijan 11.5 Bahamas 6.1 Bahrein 16.6 Bangladesh 17.4 Belgium 8.12 Bermuda 6.2 Bosnia-Herzegovina 8.20 Brazil 4.6 British Columbia 2.1 British Isles – Editorial Introduction 7.0 Brunei (Negara Brunei Darussalam) 19.4 Bulgaria 8.25 Burma (Myanmar) 17.5 California 1.4 Cambodia 19.5 Cameroon and Equatorial Guinea 13.11 Canada – Editorial Introduction 2.0 Caribbean Coast of Colombia 4.11 Caribbean Coasts, Panama to Belize 4.12 Caribbean Islands – Editorial Introduction 5.0 Caspian and Aral Seas – Editorial Introduction 11.0 Channel Islands 7.27 Chile 4.3 China 20.1 Corsica 8.13.4 Crete 8.24 Croatia 8.19 Cuba 5.1 Cumbria 7.2 Cyprus 10.2

Denmark 8.10 Dorset 7.12 Durham, Tyne and Wear 7.22 East Africa – Editorial Introduction 15.0 East Asia – Editorial Introduction 20.0 East Timor 19.9 Ecuador 4.1 Egypt (Mediterranean) 12.1 Egypt, Red Sea Coast 15.7 England and Wales 7.1 Eritrea 15.5 Essex 7.17 Estonia 8.4 Europe – Editorial Introduction 8.0 Fife 7.24.2 Fiji 23.2 Finland 8.3 Florida 1.9 Fort William to the Clyde 7.24.9 Fraserburgh to Moray Firth and Duncansby Head 7.24.4 French Guiana 4.7 Gabon, Congo, Cabinda and Zaïre 13.12 Georgia 1.10 Germany 8.9 Ghana 13.8 Gibraltar 8.14.1 Gloucestershire, Somerset and North Devon 7.8 Great Lakes (Canada) 2.7 Great Lakes (USA) 1.15 Greece 8.23 Greenland 6.3 Guinea Bissau 13.3 Gulf of St. Lawrence 2.3 Guyana 4.9 Hampshire 7.13 Hawaii 1.16 Hispaniola (Haiti and the Dominican Republic) 5.3 Iceland 6.4 India 17.2 Indian Ocean Islands – Editorial Introduction 18.0

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_List of Entries, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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List of Entries

Indian Ocean Islands 18.2 Indonesia 19.8 Iran 16.9 Iran – Caspian Sea Coast 11.1 Ireland 7.26 Isle of Wight 7.14 Isles of Scilly 7.10 Israel (with the Gaza Strip) 10.5 Italy 8.16 Ivory Coast 13.7 Jamaica 5.2 Japan 20.5 Jordan 15.8 Kaliningrad 8.7, 9.6 Kazakhstan 11.3 Kent 7.16 Kenya 15.3 Kuwait and Iraq 16.8 Lancashire, Merseyside and Cheshire 7.3 Latvia 8.5 Lebanon 10.4 Lesser Antilles 5.5 Liberia 13.6 Libya 12.2 Lincolnshire 7.20 Lithuania 8.6 Lord Howe Island – (New South Wales) 21.1.1 Louisiana 1.6 Madagascar 18.1 Malaysia – Introduction 19.2 Malta 8.17 Mauritania 13.1 Mediterranean France 8.13.3 Mexico, Caribbean Coast 4.13 Mexico, Pacific Coast 3.1 Middle East – Editorial Introduction 10.0 Mississippi 1.7 Montenegro 8.21 Morocco 12.5 Mozambique 15.1 Namibia 13.14 New Brunswick and Nova Scotia 2.2 New Caledonia and the Loyalty Islands 23.1 New South Wales 21.1 New York and New England 1.14 New Zealand 22.1

Newfoundland 2.5 Nigeria 13.10 Norfolk 7.19 Norfolk Island 21.1.2 North Africa – Editorial Introduction 12.0 North and North West Scotland (Duncansby Head to Cape Wrath and Fort William) 7.24.6 North Carolina 1.12 North Coast of France 8.13.1 North Korea 20.3 North Wales and Anglesey 7.5 Northern Canada 2.6 Northern Ireland 7.25 Northern Territory 21.3 Northumberland 7.23 Norway 8.1 Oregon 1.3 Other Atlantic Ocean Islands 6.5 Other Pacific Islands 23.4 Pacific Central America – Editorial Introduction 3.0 Pacific Coast of Colombia 3.3 Pacific Coasts of Guatemala, El Salvador, Honduras, Nicaragua, Costa Rica and Panama 3.2 Pacific Ocean Islands – Editorial Introduction 23.0 Pakistan 17.1 Papua New Guinea 19.10 Peru 4.2 Philippines 19.7 Poland 8.8 Portugal 8.15 Puerto Rico 5.4 Qatar 16.5 Queensland 21.2 Republic of Georgia 8.29 Republic of Guinea 13.4 Romania 8.26 Russian Black Sea Coast 9.2 Russian Caspian Coast 9.5, 11.4 Russian Federation – Editorial Introduction 9.0 Russian Gulf of Finland 8.3.1, 9.1 Saint-Pierre-et-Miquelon Islands 2.4 Sardinia 8.16.1 Saudi Arabia, Persian Gulf Coast 16.7 Saudi Arabia, Red Sea Coast 16.1 Scotland 7.24 Sea of Azov 8.28

List of Entries

Senegal and Gambia 13.2 Sicily 8.16.2 Sierra Leone 13.5 Singapore 19.3 Slovenia 8.18 Society Islands 23.3 Somalia and Djibouti 15.4 South Africa – Introduction 14.0 South Africa 14.1 South America – Editorial Introduction 4.0 South Asia – Editorial Introduction 17.0 South Australia 21.5 South Carolina 1.11 South Devon 7.11 South Gippsland, Wilsons Promontory and Corner Inlet 21.6.7 South Korea 20.4 South Wales 7.7 South-East Asia – Editorial Introduction 19.0 South-East Scotland 7.24.1 Southern Arabia and Oman 16.3 South-West Asia – Editorial Introduction 16.0 South-West Scotland 7.24.10 Spain 8.14 Sri Lanka 17.3 Sudan 15.6 Suffolk 7.18 Surinam 4.8 Sussex 7.15 Svalbard and Jan Mayen 8.1.1 Sweden 8.2 Syria 10.3 Taiwan 20.2 Tanzania 15.2 Tasmania 21.7 Texas 1.5 Thailand Andaman Sea Coast 19.1 Thailand: Gulf of Thailand Coast 19.3.1 The Arctic Coast of Russia 9.4 The Chatham Islands 22.2 The Inner Hebrides 7.24.7 The Isle of Man 7.4 The Netherlands 8.11

The North Coast of Cornwall 7.9.1 The Orkney and Shetland Islands 7.24.5 The Outer Hebrides 7.24.8 The Pacific Coast of Russia 9.3 The Severn Estuary 7.7.1 The South Coast of Cornwall 7.9.2 Togo and Benin 13.9 Tunisia 12.3 Turkey 10.1 Turkmenistan 11.2 Ukraine 8.27 United Arab Emirates 16.4 Uruguay 4.5 USA – Editorial Introduction 1.0 Venezuela 4.10 Victoria Introduction 21.6 Victoria: Port Phillip Bay (Point Lonsdale to Point Nepean) 21.6.4 Victoria: The East Gippsland Coast (Corner Inlet to Cape Howe) 21.6.8 Victoria: The Gippsland Lakes 21.6.9 Victoria: The Nepean Ocean Coast (Point Nepean to West Head) 21.6.5 Victoria: The Otways and South Bellarine Coast (Princetown to Point Lonsdale) 21.6.3 Victoria: The Port Campbell Coast (Mepunga to Princetown) 21.6.2 Victoria: The Western District (Discovery Bay to Mepunga) 21.6.1 Victoria: Westernport Bay, with French Island and Phillip Island 21.6.6 Vietnam 19.6 Virgin Islands 5.4.1 Washington 1.2 West Africa – Editorial Introduction 13.0 West Coast of France 8.13.2 West Wales 7.6 Western Australia 21.4 Yemen 16.2 Yorkshire 7.21

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Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_Subject Index, © Springer Science+Business Media B.V. 2010 (Dordrecht)

Subject Index

A Abandoned cliff, 208, 554, 800, 868, 872, 876, 885, 904, 905, 909, 911, 912, 940, 1003 Active dune, 41, 57, 58, 615, 621, 623, 630, 687, 879–882, 918, 1094, 1165, 1259, 1297 Aeolianite, see Dune calcarenite Arid coast, 58, 118, 207, 208, 213, 260, 342, 343, 490, 530, 531, 867, 871, 894, 897, 915, 917, 966, 971, 975, 987, 1007, 1023, 1029, 1030, 1033, 1037–1039, 1041, 1045, 1051, 1057, 1066, 1281, 1298 Artificial beach, 806, 1016, 1133, 1218 Artificial coast, 668, 728, 758–759, 1030 Asymmetrical bay, 209, 229, 896, 980, 981, 1290

B Barchan, 156, 185, 209, 214, 216, 255, 632, 645, 873, 912, 918, 919, 966, 980, 1027, 1037, 1045, 1094, 1209, 1291 Barrier, 1, 133, 183, 199, 274, 314, 351, 583, 809, 834, 863, 891, 915, 975, 989, 1021, 1057, 1087, 1115, 1190, 1225, 1404, 1425 Barrier island, 2, 156, 187, 245, 314, 352, 643, 879, 933, 1034, 1057, 1138, 1195, 1405 Barrier lagoon, 50, 164, 209, 327, 735, 738, 880, 882, 950, 983, 1126, 1457 Barrier spit, 17, 28–30, 50, 73, 77, 78, 144–146, 150, 152, 153, 156, 159, 197, 202, 221, 246, 283, 292, 324, 328, 351, 392, 451, 454, 465, 496, 547, 548, 554, 619, 620, 623, 626, 663, 683, 702, 705, 733, 776, 777, 790–792, 796, 801, 803, 835, 865, 872, 879, 936, 954, 960, 1000, 1015, 1054, 1059, 1074, 1128, 1179, 1316, 1332, 1334, 1340, 1405, 1406, 1410, 1411 Bay-mouth barrier, 114, 116, 1115 Beach drift alternations, 50, 1338 Beach erosion, 17, 56–59, 75, 83, 260, 277, 290, 410, 541, 607, 715, 724, 805, 885, 887, 888, 894, 896, 902, 944, 960, 961, 998, 1045, 1047, 1117, 1194, 1198, 1226, 1230, 1315, 1339, 1379, 1442 Beach etching, 1297 Beach mining, 307, 1099, 1143 Beach nourishment, 17, 59, 75, 83, 85, 123, 302, 351, 353, 641, 645, 669–671, 778, 805, 806, 814, 874, 952, 1131, 1144, 1198

Beach ridge, 1, 134, 186, 199, 281, 314, 380, 571, 809, 839, 859, 897, 927, 982, 993, 1022, 1060, 1094, 1111, 1195, 1232, 1418, 1428, 1465 Beach rock, 129, 201, 271, 311, 569, 839, 891, 944, 981, 989, 1023, 1067, 1097, 1114, 1242, 1423 Bevelled cliff, 369, 552 Boulder barricade, 160, 161, 171 Boulder beach, 11, 129, 165, 327, 333, 335, 336, 371, 410, 411, 422, 427, 470, 486, 496, 497, 510, 535, 541, 571, 609, 613, 616, 629, 703, 811, 818, 828, 839, 1247, 1250, 1259, 1261, 1274, 1299, 1345, 1400, 1406, 1407 Boulders, 2, 133, 185, 209, 271, 311, 350, 571, 811, 835, 909, 943, 980, 1067, 1117, 1223, 1240, 1403, 1425, 1465

C Cala, 731–733, 735–738 Calanque, 304, 378, 692, 693, 753, 755, 757, 758 Calcareous algae, 126, 280, 709, 783, 784, 1457 Calcareous sand, 125, 128, 129, 131, 204, 267, 288, 302, 311, 313, 314, 317, 325, 327, 336, 339, 344, 360, 391, 392, 398, 399, 423, 437, 463, 495, 499, 509, 511, 512, 515, 517–519, 546, 547, 549, 683, 686, 737, 781, 842, 845, 846, 853, 854, 857, 885, 887, 891, 912, 960, 1001, 1042, 1074, 1099, 1225, 1242, 1247, 1250, 1253, 1286, 1311, 1326, 1338, 1339, 1349, 1354, 1372–1374, 1377, 1413, 1428, 1430, 1434, 1456 Calcrete, 1049, 1242, 1286, 1287, 1295, 1298, 1311, 1315, 1319, 1323, 1326, 1346, 1350–1354 Canali, 761, 763 Carbonate sand, 83, 245, 290, 292–294, 311, 751, 783, 1021, 1033–1035, 1037–1039, 1041, 1066, 1071, 1268, 1354, 1379, 1457, 1458 Chalk coast, 346–348, 350, 419, 426, 428, 433, 434, 438–446, 463, 539, 540, 542, 640, 656, 658, 661, 673–675, 844 Chenier, 1, 56, 62, 65–66, 200, 221, 238, 239, 241–243, 245, 250, 352, 430, 448, 678, 792,

877, 1094, 1119, 1122, 1209, 1226, 1263, 1265, 1266, 1268, 1271–1275, 1278, 1281, 1285, 1410, 1411, 1427–1430 Chine, 143, 350, 369, 418, 421, 429, 433–435, 714 Clay cliff, 370, 423, 430, 461–463, 494, 1344, 1364, 1365 Cliff cauldron, 401, 404, 463, 517, 531, 661 Cliff sluicing, 20, 24 Coastal dune, 117, 133, 167, 170, 330, 353, 360, 366, 400, 405, 454, 458, 460, 485, 486, 547, 616, 619, 623, 666, 667, 669, 687, 713, 777, 783, 842, 896, 904, 909, 911, 919, 960, 969, 971, 975, 980, 982, 990, 1007, 1008, 1034, 1059, 1071, 1094, 1158, 1176, 1196, 1198, 1219, 1262, 1287, 1295, 1298, 1299, 1309, 1311, 1315, 1316, 1404 Coastal erratics, 146, 147, 160, 163, 221, 320, 361, 392, 404, 413, 431, 554, 607, 629, 662, 681, 1008, 1163, 1298, 1300, 1413, 1465 Coastal lagoon, 1, 2, 17, 30, 45, 109, 183, 185, 186, 210, 227, 228, 233, 266, 325, 330, 331, 583, 623, 630, 706, 725, 728, 731, 785, 789, 790, 818, 899, 922, 935, 940, 948, 975, 984, 990, 1037, 1041, 1052, 1069, 1071, 1075, 1094, 1119, 1160, 1167, 1205, 1225, 1233, 1234, 1286, 1288, 1315, 1316, 1377, 1385, 1398, 1417 Coastal landslide, 350, 506, 1332, 1415 Coastal plain, 1, 136, 183, 199, 274, 320, 350, 573, 809, 835, 861, 897, 915, 982, 987, 1021, 1057, 1087, 1111, 1187, 1225, 1430 Coastal reclamation, 20–21, 24, 48, 50, 141–144, 243, 260, 262, 348, 353, 358–359, 382–383, 385, 389, 431, 438, 441, 443–447, 450, 457, 459, 461, 479, 536–538, 546, 554, 563, 641, 645, 656, 666, 669–670, 676, 686, 711–712, 720, 725, 743, 779, 925, 957, 1038, 1041–1042, 1111, 1117, 1122, 1125, 1127, 1128, 1131–1133, 1152, 1161, 1191, 1201, 1208, 1217–1218, 1314, 1359, 1364, 1376 Coastal terrace, 46, 47, 71, 145, 213, 237, 415, 471, 475, 486, 490, 491, 529–531, 541, 551, 736, 827, 909, 950, 1001, 1031, 1054, 1062, 1117, 1121, 1127, 1201, 1207, 1321, 1322, 1403, 1405 Coastal tor, 400, 402, 414–416, 565, 1130

Eric C.F. Bird (ed.), Encyclopedia of the World’s Coastal Landforms, DOI 10.1007/ 978-1-4020-8639-7_Subject Index, © Springer Science+Business Media B.V. 2010 (Dordrecht)

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Subject Index

Coastal waterfall, 335, 350, 394, 1415 Continental shelf, 1, 154, 183, 199, 274, 340, 535, 687, 811, 833, 896, 918, 989, 1021, 1059, 1089, 1144, 1201, 1225, 1403, 1465 Coral reef, 2, 183, 199, 271, 311, 954, 975, 987, 1021, 1057, 1087, 1111, 1187, 1225, 1423 Cuspate foreland, 2, 79, 81, 85–87, 94, 106, 109, 110, 138, 209, 232, 236, 328, 353, 371, 443, 471, 491, 505, 640, 653, 656, 717, 779, 900, 1046, 1049, 1096, 1137, 1165, 1177, 1178, 1181, 1185, 1266, 1271, 1272, 1286, 1287, 1298, 1302, 1340, 1341, 1375, 1384, 1429, 1432, 1433, 1466 Cyclone impacts, 1255, 1258, 1268, 1279, 1284–1285

D Dalmatian coast, 761–765, 1081 Delta cycle, 62 Deltaic plain, 25, 61–67, 195, 197, 228, 691, 706, 722–724, 733, 774, 778, 792, 796, 831, 838, 839, 877, 878, 886, 894, 900, 944, 1052, 1077–1079, 1081, 1084, 1111, 1135, 1152, 1154, 1163–1165, 1173, 1178–1180, 1189, 1191, 1229, 1340, 1383, 1430, 1442 Delta sediments, 91, 826, 879, 886, 944 Doline, 314, 724, 753, 757, 1264, 1295, 1296 Driftwood, 23, 1392 Drumlins, 2, 113, 133, 141, 151–153, 355, 367, 368, 537, 538, 545, 548–551 Dune calcarenite, 3, 229, 289, 314, 397, 695, 831, 891, 915, 980, 989, 1037, 1094, 1225, 1413, 1426

F Fiard, 572, 828 Fiord, 1, 12, 135, 138, 170, 172, 173, 574, 577, 652, 827 Fjördur, 324, 326, 331 Flandrian marine transgression, 44, 348, 350, 351, 354, 358, 384, 430, 545, 546, 674, 680, 753, 922, 931, 1066, 1089, 1096 Flysch, 812, 1412 Font, 1297

G Geo, 33, 86, 105, 300, 347, 486, 492, 500, 547, 548, 582, 695, 785, 894, 901, 907, 929, 1065, 1143, 1189, 1395 Glaciated coast, 1, 10, 12, 113, 500, 523, 538, 1465, 1467 Granite coast, 400, 402, 731, 733, 1124 Gravel beach, 10, 18, 118, 133, 150, 172, 173, 199, 219, 222, 286, 304, 328, 336, 351, 509, 609, 656, 663, 681, 801, 805, 815, 842, 852, 969, 979, 1030, 1085, 1100, 1114, 1153, 1155, 1217, 1339, 1407, 1415, 1467 Guano, 20, 204, 209, 335, 971, 972, 1107, 1314, 1375, 1420, 1456, 1465

M H Honeycomb weathering, 21, 1275, 1352, 1384 Hoodoo, 1198 Hurricane impacts, 55, 57, 59, 63, 75, 79, 152, 266, 267, 279, 299, 302, 303, 307, 616, 1059

I E Earthquake effects, 3, 12, 33–35, 44, 128, 188, 194, 195, 197, 207, 395, 401, 778, 1008, 1059, 1071, 1081, 1108, 1114, 1118, 1158, 1183, 1408, 1409 Earthquakes, 6, 11, 16, 48, 183, 193–195, 199, 227, 753, 781, 783, 785, 809, 833, 891, 909, 957, 1054, 1057, 1064, 1151, 1157–1159, 1169, 1175, 1176, 1187, 1215, 1219, 1220, 1307, 1403, 1413 Emerged beach, 172, 173, 221, 348, 503, 510, 523, 524, 538, 545, 548, 591, 917, 933, 975, 1029, 1032, 1072, 1074, 1081, 1099, 1187, 1300, 1465 Emerged coast, 18, 161, 287, 320, 327, 475, 486, 490, 491, 511, 525, 552, 827, 866, 975, 1003, 1030, 1038, 1045, 1048, 1071, 1081, 1135, 1147, 1222, 1440, 1446, 1448 Emerging coast, 34, 67, 512, 619, 623, 1198 Eolianite, see Dune calcarenite Escarpment cliff, 356, 357, 428, 444, 463, 714, 716, 1239, 1240, 1243 Esker, 113, 141, 469, 491, 518, 545, 589, 590, 595, 609, 1465 Estuarine meander, 108 Evaporite deposits, 208, 889, 1281, 1291 Exfoliation, 597, 695, 1134, 1374

Land reclamation, 48, 141, 262, 431, 438, 444, 459, 536–538, 563, 686, 711, 957, 1038, 1111, 1117, 1121, 1122, 1124, 1125, 1127, 1128, 1152, 1191, 1346 Landslide, 1, 163, 194, 205, 323, 350, 585, 809, 857, 873, 1032, 1151, 1239, 1415 Lavaka, 229 Lava lobe, 131 Liman (lagoon), 787, 796, 797, 801–803, 811, 812, 819, 826, 827 Limestone, 99, 159, 183, 252, 271, 311, 345, 581, 809, 831, 868, 891, 960, 1000, 1001, 1021, 1060, 1094, 1111, 1196, 1279, 1413, 1423 Limestone cliff, 304, 696, 701, 746, 765, 776, 1031, 1032, 1168 Limestone coast, 178, 377, 728, 761–763, 767, 1447 Litorina transgression, 605, 609, 614, 639, 649, 654 Littoral cell, 33, 36, 37, 39, 40, 50, 127, 129, 178, 893 Longshore drift, 27, 181, 254, 324, 443, 592, 615, 628, 800, 819, 855, 1466

Ice cliff, 217, 319, 320, 336, 337, 468, 1463, 1464, 1466, 1468 Ice coast, 169, 171, 172, 583, 1464 Isostatic movements, 16, 142, 348, 475, 606, 615, 628

K Karst, 85, 127, 134, 169, 175, 176, 266, 269, 273, 286, 289, 290, 293, 303, 314, 317, 320, 445, 550, 712, 714, 716, 717, 724, 741, 753–755, 761–765, 769, 774, 783, 826, 838, 850, 991, 993, 999, 1000, 1042, 1084, 1089, 1092, 1111, 1113, 1114, 1119, 1121, 1136, 1149, 1156, 1196, 1242, 1244, 1248, 1250, 1264, 1296, 1319, 1353, 1418, 1431, 1434, 1458 Klint, 601, 605–607, 656, 658, 661, 714

L Lagoons, 1, 133, 183, 199, 271, 314, 351, 583, 809, 834, 863, 891, 915, 975, 987, 1023, 1057, 1087, 1119, 1190, 1225, 1405, 1423 Lake level fluctuations, 121–123, 177, 180, 792, 1388 Lakes, 2, 133, 185, 221, 305, 317, 347, 589, 823, 833, 894, 917, 975, 990, 1029, 1057, 1117, 1205, 1225, 1403, 1430, 1467

Machair, 491, 495, 496, 508, 510–512, 515, 516, 518–522, 549 Managed retreat, 353, 385, 449, 1125 Mangroves, 2, 183, 199, 271, 311, 915, 975, 987, 1021, 1057, 1089, 1111, 1187, 1228, 1418, 1423 Megacliff, 214, 547–548, 551, 1456, 1466 Megaripples, 147, 148, 205, 392, 674, 678 Megatidal coast, 134, 142 Mélange, 1, 15, 16, 29, 45, 935 Mining impacts, 48, 274, 351, 354, 399, 401, 406, 407, 429, 466, 857, 1084, 1089, 1099, 1125, 1143, 1161, 1185, 1398, 1423, 1425, 1428, 1432, 1433 Morainic coast, 579 Mound, 126, 141, 330, 385, 455, 461, 531, 645, 826, 877, 878, 943, 948, 1015, 1060, 1083, 1160, 1176, 1403, 1407, 1408, 1414, 1465 Mudbank, 240–243, 245, 387, 408, 441, 475, 1059, 1116

N Natural arch, 15, 49, 159, 206, 317, 327–328, 331, 335, 350, 365, 373, 377, 467, 471, 490, 495, 496, 499, 507, 509, 516–517, 520, 548, 550, 675, 677, 753–754, 982, 1171, 1252, 1284, 1320, 1323, 1369

O Ocean swell, 1, 141, 183, 199, 271, 311, 397, 569, 815, 907, 915, 976, 987, 1029, 1057, 1087, 1111, 1187, 1225, 1413 Old Hat island, 27, 28, 1243

Subject Index

P Peat, 99, 156, 163, 165, 221, 229, 243, 246, 281, 311, 323, 339, 354, 372, 448, 449, 459, 494, 496, 503, 512, 520, 542, 543, 546, 549, 561, 576, 619, 632, 641, 645, 655, 666, 669, 670, 687, 810, 820, 1119, 1124, 1126, 1138, 1142, 1208, 1210, 1389, 1413, 1415, 1416, 1420, 1463, 1467 Pebbles, 2, 159, 207, 281, 328, 351, 571, 813, 835, 907, 943, 1032, 1070, 1117, 1189, 1240, 1419, 1431, 1465 Periglacial effects, 398, 824, 826 Periglaciated coast, 348, 350, 361, 398, 420, 1467 Permafrost, 8, 10, 133, 134, 169, 174, 176, 582, 809, 810, 823, 824, 1463, 1465 Permatang (beach ridges in Malaysia), 1112, 1119, 1122 Pindan, 1281, 1283, 1285 Playa, 201, 209–211, 229, 251, 294, 702–705, 973 Pleistocene glaciations, 15, 133, 167, 177, 199, 345, 346, 353, 545, 569, 809, 1070 Progradation, 17, 54–56, 58, 59, 62–66, 73, 85, 89, 93, 102, 122, 205, 228, 236, 241, 250, 254, 274, 360, 483, 487, 607, 642, 654, 656, 670, 686, 706, 773, 774, 776, 803, 838, 877, 900, 919, 927, 928, 930, 944, 949, 952, 969, 992, 1003, 1034, 1042, 1066, 1069, 1070, 1084, 1119, 1127, 1149, 1160, 1165, 1167, 1176–1178, 1191, 1197, 1209, 1217, 1226, 1243, 1249, 1261, 1264, 1274, 1362, 1374, 1377, 1379, 1380, 1404, 1409, 1410, 1425, 1428, 1437, 1442, 1445 Prograded beach, 233, 236, 254, 453, 707, 837, 1152, 1155, 1158, 1169, 1218, 1330, 1361 Pumice, 267, 337, 780, 853, 1161, 1456, 1468

R Radiocarbon dates, 194, 1003, 1126, 1237, 1249, 1444, 1458, 1460 Raised beach, see Emerged beach Rasa, 223, 225, 699, 701, 907, 909 Ria, 48, 107, 119, 133, 156, 161, 199, 200, 213, 216, 217, 232, 251, 275, 346, 348, 376, 401, 408, 413–415, 433, 546, 550, 551, 553, 554, 590, 678, 680, 681, 683, 689, 701, 702, 704, 705, 713, 731, 753, 754, 762, 764, 766, 767, 776, 777, 796, 797, 800, 815, 816, 823, 927, 929, 931, 933, 948, 1032, 1067, 1092, 1095, 1096, 1128, 1161, 1179, 1187, 1197, 1201, 1204, 1207, 1208, 1213, 1215, 1216, 1219, 1222, 1235, 1281, 1361, 1395, 1403, 1407, 1408, 1425, 1430, 1433 Ridge and runnel, 360, 536, 1123, 1124

Salt marsh, 2, 90, 141, 199, 311, 352, 355, 571, 873, 904, 927, 982, 1027, 1065, 1201, 1305, 1428 Salt pan, 380, 742, 743, 754, 766, 812, 893, 919, 969–972, 1135, 1256, 1260, 1262, 1266 Scar, 381, 420, 421, 433, 531, 1240, 1321 Sea ice, 8–11, 134, 142, 144, 161, 167, 169, 172–174, 311, 319, 337, 340, 369, 375, 411, 582, 584, 587, 639, 681, 809, 823, 1213, 1464, 1465, 1467 Sea level changes, Aral Sea, 810, 859, 889–890 Sea level changes, Caspian Sea, 859, 861–869, 871–880, 882, 885–887, 890 Sea wall, 18, 223, 305, 328, 353, 670, 842, 887, 1047, 1067, 1098, 1160, 1193, 1226 Sebkha, 250, 252, 256, 893, 896–898, 907, 912, 913, 915, 919, 1007, 1013, 1015–1017, 1021, 1022, 1027, 1033–1036, 1038, 1041, 1042, 1045–1047, 1051, 1281, 1282 Segmentation, 11, 62, 1074, 1390 Sharm, 1013, 1021, 1023, 1024, 1030, 1038, 1041 Shelly beach, 223, 343, 447, 449, 479, 505, 555, 705, 872, 881, 1003, 1004, 1015, 1115, 1121, 1125, 1274, 1276, 1313, 1337, 1339, 1349, 1360, 1369, 1375, 1418 Shingle, 2, 143, 194, 319, 348, 583, 813, 834, 891, 933, 1015, 1083, 1164, 1255 Shingle beach, 29, 144, 207, 282, 346, 571, 809, 835, 905, 1231, 1467 Shoals, 18, 151, 197, 201, 283, 311, 351, 644, 809, 853, 867, 918, 990, 1022, 1057, 1119, 1194, 1234, 1417, 1428 Shore ice, 133, 134, 141, 142, 153, 155, 159–161, 167, 323, 597, 606, 607, 615, 619, 623, 641, 663, 809, 1187 Shoreward drifting, 50, 107, 207, 467, 526, 867, 871, 1072, 1226, 1379, 1380 Skerry, 573, 575, 827, 828 Slumping cliff, 18, 250, 422, 449 Spartina, 49, 147, 152, 156, 157, 161, 221, 227, 228, 353, 354, 385, 387, 388, 394, 417, 429, 430, 433, 438, 444, 533, 716, 982, 984, 1194, 1228, 1365, 1366, 1371, 1376, 1398, 1399 Stack, 1, 141, 194, 204, 306, 311, 350, 571, 817, 834, 924, 1003, 1032, 1057, 1114, 1216, 1241, 1413, 1431, 1464 Steep coast, 1, 133, 189, 209, 311, 350, 583, 809, 831, 905, 940, 980, 1016, 1027, 1057, 1095, 1111, 1187, 1259, 1420, 1427 Storm surge, 38, 151, 214, 300, 345, 601, 809, 834, 861, 952, 1021, 1057, 1118, 1208, 1225, 1437 Storm surge effects, 868 Strandflat, 136, 324, 576 Subsiding coast, 43, 99, 1440 Swampy shore, 227, 238, 255, 266, 925, 991, 1001, 1074, 1136, 1170

S Salt dome, 646, 1022, 1027, 1034, 1035, 1051, 1052 Salt dome island, 1034, 1035, 1038

T Tafoni, 28, 733, 734, 966, 1299 Tanne, 925, 927, 929

Tectonic movements, 44, 189, 195, 254, 260, 601, 606, 615, 699, 723, 775, 781, 885, 894, 987, 995, 1011, 1013, 1045, 1081, 1111, 1157, 1159, 1217, 1219, 1227, 1305, 1359, 1377, 1413 Terrane, 1294 Tidal bore, 141, 142, 148, 348, 676, 1064, 1081, 1118, 1189, 1193, 1268, 1278, 1281 Tidal current scour, 13, 369, 1057, 1146, 1176, 1379 Tidal delta, 73, 85, 90, 93, 159, 206, 728, 1035, 1269, 1347, 1405 Tidal divide, 18, 1360, 1362, 1364, 1377 Tidal flat, 358, 669 Tidal inlet, 57, 64, 86, 89, 90, 93, 94, 98, 106, 110, 113, 156, 159, 193, 233, 330, 378, 644, 666, 728, 927–930, 1030, 1057, 1092, 1146, 1176, 1177, 1376, 1379, 1405 Tidal mudflat, 378, 1193 Tidal waterfall, 1281 Tide-dominated coast, 89, 93, 142, 145, 195 Tombolo, 18, 145, 206, 281, 331, 410, 585, 817, 844, 904, 935, 1027, 1059, 1115, 1194, 1234, 1410, 1425 Transgressive dune, 17, 89, 213, 214, 992, 1074, 1225, 1311, 1326 Trottoir, 709, 849, 851, 852 Tsunami, 1, 3, 12, 16, 29, 33, 44, 45, 188, 193, 195, 197, 403, 496, 553, 554, 778, 780, 783, 809, 909, 1008, 1059, 1065, 1070, 1071, 1108, 1114, 1115, 1118, 1122, 1151, 1158–1161, 1163, 1164, 1176, 1183, 1215, 1216, 1222, 1223, 1230, 1235, 1238, 1413, 1441, 1457 Tufa towers, 320

V Volcanic coast, 126, 131, 289, 817, 1161, 1221, 1456, 1457 Volcanic islands, 188, 189, 206, 271, 294, 311, 323, 328, 336, 338, 340, 342, 344, 780, 781, 955, 1083, 1087, 1098, 1100, 1102, 1161, 1169, 1180, 1185, 1222–1224, 1423, 1434, 1445, 1449, 1451, 1455–1457, 1467, 1468 Volcanoes, 34, 125, 128, 129, 131, 206, 250, 260, 271, 307, 311, 324, 333, 337–339, 342, 803, 817, 853, 876, 1021, 1057, 1061, 1083, 1111, 1151, 1157, 1158, 1162, 1167, 1169, 1172, 1175, 1179, 1180, 1184, 1215, 1222, 1423, 1441, 1443, 1446, 1447, 1451, 1452, 1455, 1456

W Weathering, 16, 144, 201, 290, 317, 351, 569, 815, 835, 904, 917, 993, 1042, 1129, 1191, 1227, 1403, 1429 Wetland, 43, 62–64, 66, 71, 77, 89, 113, 122, 123, 129, 175, 221, 225, 279, 281, 300, 304–306, 308, 385, 389, 538, 616, 674, 705, 742, 743, 778, 877, 1119, 1127, 1159, 1196, 1249, 1288, 1290, 1398

Z Zawn, 365, 401

1493