The Physical Geography of the Mediterranean (Oxford Regional Environments)

The Physical Geography of the Mediterranean

THE OXFORD REGIONAL ENVIRONMENTS SERIES PUBLISHED

The Physical Geography of Africa edited by William M. Adams, Andrew S. Goudie, and Antony R. Orme The Physical Geography of North America edited by Antony R. Orme The Physical Geography of Northern Eurasia edited by Maria Shahgedanova The Physical Geography of Southeast Asia edited by Avijit Gupta The Physical Geography of Fennoscandia edited by Matti Seppälä The Physical Geography of Western Europe edited by Eduard A. Koster The Physical Geography of South America edited by Thomas T. Veblen, Kenneth R. Young, and Antony R. Orme FORTHCOMING

The Physical Geography of the British Isles edited by Adrian Parker

The Physical Geography of the Mediterranean edited by Jamie Woodward

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Great Clarendon Street, Oxford OX2 6DP Oxford University Press is a department of the University of Oxford. It furthers the University’s objective of excellence in research, scholarship, and education by publishing worldwide in Oxford New York Auckland Cape Town Dar es Salaam Hong Kong Karachi Kuala Lumpur Madrid Melbourne Mexico City Nairobi New Delhi Shanghai Taipei Toronto With offices in Argentina Austria Brazil Chile Czech Republic France Greece Guatemala Hungary Italy Japan Poland Portugal Singapore South Korea Switzerland Thailand Turkey Ukraine Vietnam Oxford is a registered trademark of Oxford University Press in the UK and in certain other countries Published in the United States by Oxford University Press Inc., New York © The several contributors 2009 The moral rights of the author have been asserted Database right Oxford University Press (maker) First published 2009 All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted, in any form or by any means, without the prior permission in writing of Oxford University Press, or as expressly permitted by law, or under terms agreed with the appropriate reprographics rights organization. Enquiries concerning reproduction outside the scope of the above should be sent to the Rights Department, Oxford University Press, at the address above You must not circulate this book in any other binding or cover and you must impose the same condition on any acquirer British Library Cataloguing in Publication Data Data available Library of Congress Cataloging in Publication Data The physical geography of the Mediterranean / edited by Jamie Woodward. p. cm. – (Oxford regional environments series) ISBN 978–0–19–926803–0 1. Physical geography–Mediterranean Region. 2. Mediterranean Region–Geography. I. Woodward, Jamie C. GB178.P48 2009 508.3182’2–dc22 2009005674 Typeset by SPI Publisher Services, Pondicherry, India Printed in Great Britain on acid-free paper by CPI Antony Rowe, Chippenham, Wiltshire ISBN 978–0–19–926803–0 1 3 5 7 9 10 8 6 4 2

Frontispiece. A reconstruction of ancient Sparta on the alluvial plain of the Evrotas River in southern Greece. The view looks to the west to the peaks of Mount Taygetos. Reproduced with permission from an illustration by J. P. Mahaffy published in 1890.

‘. . . the Mediterranean can’t be reduced down to one landscape or one lifestyle. Could there be a more tenuous link than that which bonds, upon its shores, the luxuriant landscapes of the coast and the arid deserts inland? . . . The Mediterranean lover can recognise the imperceptible variations that alter the texture of a valley, the hue of a city, and the quite special light of a particular bay. But our ideas about the Mediterranean are stubborn. Its diversity doesn’t stop us from seeing a Mediterranean nature, climate and landscape at work all over.’ (Girard, 2001, p. 36) Girard, X. (2001), Mediterranean: from Homer to Picasso. Assouline, New York.

For Sam and Alex

Foreword The Physical Geography of the Mediterranean is the eighth in a series of advanced books that is being published by Oxford University Press under the rubric of Oxford Regional Environments. The aim of the series is to provide a durable statement of physical conditions on each of the continents, or major regions within those continents. Each volume includes a discussion of the systematic framework of the region (for instance, tectonism, climate, biogeography), followed by an evaluation of dominant environments (such as mountains, forests, and deserts) and their linkages, and concludes with a consideration of the main environmental issues related to the human use and misuse of the land (such as resource exploitation, agricultural and urban impacts, pollution, and nature conservation). While books in the series are framed within an agreed context, individual books seek to emphasize the distinctive qualities of each region. We hope that this approach will provide a coherent and informative basis for physical geography and related sciences, and that each volume will be an important and useful reference source for those concerned with understanding the varied environments of the continents. Andrew Goudie, University of Oxford Antony Orme, University of California, Los Angeles

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Preface and Acknowledgements Scholars have been fascinated by the Mediterranean environment since classical times. The Mediterranean, including the sea itself and the climates, landscapes, and ecosystems around it, forms one of the most intensively studied natural laboratories on Earth and recent years have seen remarkable advances in our understanding of many components of its physical geography. These advances have come from the efforts of countless researchers in many disciplines. This book explores the evolution and functioning of the Mediterranean environment by drawing upon findings derived from studies of modern processes as well as data from long records of change. It assesses the main drivers of environmental change (including tectonic processes and climate change) and their effects— with particular emphasis on the glacial-interglacial cycles of the Quaternary ice age. Several chapters also draw upon sedimentary, archaeological and historical records to examine the nature of human-environment interactions during the postglacial period in order to explore the role of humans in shaping the landscapes and habitats we see today. This book examines these processes and the key debates, and places natural hazards and current environmental issues in long term context. The physical geography of the Mediterranean involves a big canvas and a vast literature. This book is the first modern synthesis of the physical geography of the entire Mediterranean region and it was conceived, from the outset, as a multi-authored volume and as an international, multidisciplinary team effort. It incorporates the talents and experience of thirty five scholars who, between them, have direct field experience in all of the countries that border the Mediterranean Sea and in all of the major islands and marine basins within it. Many of the contributors are the leaders in their fields. This body of expertise gives this volume an authority that could not be attained by a singleauthored text. The twenty-three chapters are organised into four main parts—each with an Editorial Introduction that draws out major themes—under the following headings: I. II. III. IV.

The Physical and Biological Framework (Chapters 1 to 5) Process and Change in Specific Environments (Chapters 6 to 14) Hazards (Chapters 15 to 19) Environmental Issues in the 21st Century (Chapters 20 to 23)

This book will appeal to all scholars and students of geography, Earth science and ecology who are involved in the study of the Mediterranean and all who are interested, more broadly, in Mediterranean environments, geomorphology, natural hazards, Quaternary environmental change, biodiversity and conservation, and the human impact on the natural environment. The last decade has seen the publication of several excellent books on the history and prehistory of the Mediterranean world and something of a revival in considering the entire region as a unit of study. Research into the human past in the Mediterranean cannot be separated from an understanding of the opportunities and constraints offered by such a dynamic (and often hazardous) and varied environment. I hope, therefore, that this book will also have wide appeal to archaeologists and historians of the Mediterranean world and to all who are interested, more generally, in humanenvironment interactions—especially over extended timescales Interacting with so many authors and disciplines has been a challenging but hugely rewarding experience and I would like to offer my warmest thanks to all the authors for their contributions and for their patience and cooperation during the reviewing process and the final editorial and production stages. Before final editing, all

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twenty-three chapters were formally reviewed by the editor and by at least one external reviewer and I would like to thank the following for their generous help in this task: Clive Agnew (Manchester), Harriet Allen (Cambridge), Nick Ambraseys (London), Grant Bigg (Sheffield), Louise Bracken (Durham), Rob Bryant (Sheffield), David Chester (Liverpool), Jacques-Louis de Beaulieu (Marseilles), Simon Davis (Lisbon), Mick Frogley (Sussex), Dick Grove (Cambridge), Philip Hughes (Manchester), Ian Lawson (Leeds), Mike Leeder (Norwich), John Lewin (Aberystwyth), Sarah Lindley (Manchester), Donatella Magri (Milan), Allen Perry (Swansea), Rick Shakesby (Swansea), Heather Viles (Oxford), Des Walling (Exeter), Tony Waltham (Nottingham), and Kathy Willis (Oxford). Jeff Blackford, Karen Exell, Philip Hughes, John Lewin, and Chris Perkins also provided helpful comments on the four Editorial Introductions. In the early stages of this project the series editors, Andrew Goudie and Antony Orme, provided very helpful feedback on the initial book proposal. I am also grateful for the support of my colleagues at the University of Manchester and to Tim Allott (Head of Geography) who provided funding to cover the costs of translating Chapter 21. Avi Gupta provided wise words of support—especially in the early stages when his volume, The Physical Geography of Southeast Asia, was nearing completion. I am also very grateful to Anne Ashby, Dominic Byatt, Lizzy Suffling, and Louise Sprake at Oxford University Press who have provided valuable advice and encouragement at various stages. I would also like to thank Sylvie Jaffrey and Debbie Sutcliffe for all their help at the copy-editing and proof checking stages. Of the many people who have helped along the way getting this book to publication, I would especially like to thank Nick Scarle from the Cartographic Unit at the University of Manchester. Nick managed the figure and photograph database for this book and redrew most of the maps and figures from scratch. This was a huge undertaking, but as the figures and photographs arrived in Manchester from distant lands via email, CD, and hard copy (at least two were barely decipherable scans of sketches made on paper napkins) Nick prepared them for publication with his trademark expertise and endless patience and good humour. I would also like to thank Graham Bowden in the Cartographic Unit at Manchester who worked on the figures in Chapters 13 and 17. I would also like to thank John Prag at the Manchester Museum who provided the source book for the frontispiece. My own research in the Mediterranean began in 1986 during my Ph.D. research in Cambridge on the Klithi Project in north-west Greece directed by Geoff Bailey (Bailey, 1997). I was fortunate at that time to cut my teeth in the field with Mark Macklin and John Lewin and our collaborations have continued ever since. I am delighted that they were able to co-author chapters with me in the current volume. I must also thank all the undergraduate students at Leeds and Manchester who either took one of the various incarnations of my final year course on Mediterranean Quaternary Environments or participated in the field course to south-east Spain—their comments and feedback have helped to shape this book. I am also extremely grateful to the following research students who have participated in Mediterranean adventures: George Christopolos, Graham Smith, Rob Hamlin, Suzanne Hewitt, Maroulia Zorzou, Philip Hughes, Mike Morley, and Rose Wilkinson. My research in the Mediterranean has involved collaboration with many projects and individuals and I have been fortunate to receive support and encouragement from many colleagues. I would especially like to thank the following: Geoff Bailey, Graeme Barker, Alex Chepstow-Lusty, Ian Foster, Mick Frogley, Clive Gamble, Philip Gibbard, David Gilbertson, Paul Goldberg, Dick Grove, Philip Hughes, Takis Karkanas, Mike Kirkby, Eleni Kotjabopoulou, Mike Krom, Henry Lamb, John Lewin, Mark Macklin, Rolfe Mandel, Mark Pluciennik, Jim Rose, Nick Shackleton, John Thornes, Charles Turner, Chronis Tzedakis, Claudio Vita-Finzi, Richard West, Bob Whallon, and Martin Williams. I would also like to pay a special tribute to John Thornes who died in the summer of 2008 when this book was in the final stages of production. John was an inspirational

Preface and Acknowledgements

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figure in physical geography who was passionate about the Mediterranean, its people, and its geomorphology. For many years John coordinated and led the MEDALUS projects funded by the European Union and his research teams produced many key insights into land degradation and river basin processes in the semi-arid Mediterranean. John wrote two chapters for this book (one of them co-authored with me and Francisco LópezBermúdez) and it was a great pleasure to be able to work closely with him throughout. He was a stimulating, supportive, and generous collaborator and I hope that he would have been pleased with this book. Finally I must thank Jenny, Sam, and Alex for their love and support. J. W. Manchester November 2008

Reference Bailey, G. N. (ed.) (1997), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece (2 vols.) McDonald Institute for Archaeological Research, Cambridge. The authors, editor, and publisher thank the following who have kindly given permission for the modification and use of copyright material: Frontispiece: ‘Ancient Sparta’ in Greek Pictures, p. 173, J.P. Mahaffy (1890) published by the Religious Tract Society. Reproduced with permission from Lutterworth Press. Fig. 1.7: from Junta de Andalucía Ortofotografía digital de Andalucía (ISBN 84-95083-96-5) with permission. Figs. 3.3(a) and (b): from Futures for the Mediterranean Basin: The Blue Plan, 6–7, Grenon and Batisse (1989) with permission from Oxford University Press. Table 3.4: from http://natural-hazards.jrc.ec.europa.eu/activities_flood_flashflood.html with permission from Jutta Thielen-del Pozo at JRC European Commission. Figs. 5.1, 5.3, 5.6, 5.11, 5.12, 5.15, 5.17, 5.20 and Table 5.1: from Biology and Wildlife of the Mediterranean Region, Blondel and Aronson (eds.) (1999) with permission from Oxford University Press. Fig. 6.3: from Catena 40, 3–17, ‘The effect of land parameters on vegetation performance and degree of erosion under Mediterranean conditions’, Kosmas et al. (2000) with permission from Elsevier. Fig. 6.4: from Catena, 28, 157–169, ‘Soils in the Mediterranean region: what makes them different?’ Yaalon (1997) with permission from Elsevier. Fig. 6.9: from Soil and Tillage Research 85, 123–142, Assessment of tillage erosion by mouldboard plough in Tuscany (Italy)’, De Alba et al. (2006) with permission from Elsevier. Fig. 6.11: from Geomorphology 13, 87–99, ‘Short and long term effects of bioturbation on soil erosion, water resources and soil development in an arid environment’, Yair (1995) with permission from Elsevier. Fig. 6.12: from ‘Mechanisms of overland flow generation and sediment production on loamy and sandy soils with and without rock fragments’, Poesen et al. in Overland Flow Hydraulics and Erosion Mechanics, Parsons and Abrahams (eds.) (1992) with kind permission from Professor Jean Poesen. Fig. 6.13: Reproduced with kind permission from Juan Puigdefábregas. Fig. 6.16: from Geomorphology 26, 239–251, ‘Factors underlying piping in the Basilicata region, southern Italy’, Farifteh and Soeters (1999) with permission from Elsevier. Fig. 6.18: from Geoderma 105, 125–140, ‘Soil erosion caused by extreme rainfall events: mapping and quantification in agricultural plots from very detailed digital elevation models’, Martínez Casasnovas et al. (2002) with permission from Elsevier. Fig 6.20: Earth and Planetary Science Letters 195, 169–183, ‘Power-law correlations of landslide areas in central Italy’, Guzzetti et al. (2002) with kind permission from Fausto Guzzetti and Elsevier.

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Preface and Acknowledgements Fig. 6.22: from Engineering Geology 70, 109–130, ‘Instability conditions of marly hillslopes: towards landsliding or gullying? The case of the Barcelonnette Basin, south east France’, Maquaire et al. (2003) with permission from Elsevier. Fig. 6.23: from Engineering Geology 58, 89–107, ‘Landslide fatalities and the evaluation of landslide risk in Italy’, Guzzetti (2000) with permission from Elsevier. Fig. 6.24: from Bulletin of Engineering Geology and the Environment 59, 87–97, ‘History of the 1963 Vaiont slide: the importance of geological factors’, Semenza and Ghirotti (2000) with permission from Elsevier. Fig 7.2: from Israel Journal of Botany 39, 481–508, ‘Global change: vegetation, ecosystems, and land use in the southern Mediterranean basin by the mid 21st century’, Le Houerou (1990) with permission from LPP Ltd. Fig 7.5: from ‘The study of plant groupings in the countries surrounding the Mediterranean: some methodological aspects’, Quézel, 87–93, in Ecosystems of the World, vol. 11 Mediterranean-type Shrublands, di Castri et al. (eds.) (1981) with permission from Elsevier. Fig 7.6: from materials in: (1) the Atlas d’Aréologie Périméditerranéenne Daget, (1980) (Maison d’edition: Institut de Botanique, Montpelier) with permission from the author; (2) ‘Definition of the Mediterranean region and the origin of its flora’, Quézel, 8–24, in Plant Conservation in the Mediterranean Area, in Gómez-Campo (ed.) (1985) with permission from Springer (Kluwer); (3) The Holocene 1, 157–161, ‘The recent distribution of Pinus brutia: a reassessment based on dendroarchaeological and dendrohistorical evidence from Israel’ Biger and Liphschitz (1991) with permission from Hodder Headline. Fig 7.13: from The Holocene, p.190 Roberts (1998) with permission from Blackwell. Fig 8.2: from Futures for the Mediterranean Basin: The Blue Plan, p. 28, Grenon and Batisse (1989) with permission from Oxford University Press. Fig 8.3: from Futures for the Mediterranean Basin: The Blue Plan, p. 221, Grenon and Batisse (1989) with permission from Oxford University Press. Fig. 8.8: from Geoderma, 15, 61–70, ‘Vegetal cover to estimate soil erosion hazard in Rhodesia’, Elwell and Stocking (1976) with permission from Elsevier Science Table 8.1 from Journal of the Geological Society, London 162, 879–908, ‘Recent evolution of a Mediterranean deltaic coastal zone: human impacts on the Inner Thermaikos Gulf’, Kapsimalis et al. (2005) with permission from The Geological Society of London. Fig. 8.16: from ‘Erosion and sediment yield in mountain areas of the world’ Dedkov and Moszherin, 29–36, in Erosion, Debris Flows and Environment in Mountain Regions of the World, Walling et al. (eds.) IAHS Publication No. 209 (1992) with permission from The International Association of Hydrological Sciences Press. Fig. 10.12: from Catena Supplement 25, p. 88, ‘Environmental change and human impacts on the Mediterranean karsts of France, Italy and the Dinaric region’, Gams et al. (1993) with permission from Catena Verlag. Fig. 10.13(a) and (c): from Zeitschrift für Geomorphologie 22, 170–81, ‘The Polje: The problems of its definition’, Gams (1978) with permission from Gebrüder Borntraeger Science Publishers (BerlinStuttgart). Figs. 10.15, 10.16(a) and (c): with permission from Dr Tony Waltham. Fig 10.20: from Catena Supplement 25, p. 8, ‘Environmental change and human impact on karst terrains’, Williams (1993) with permission from Catena Verlag. Fig 11.2(a): from Proceedings of the Prehistoric Society, 32, 1-29, ‘The climate, environment and industries of Stone Age Greece: Part II’, Higgs and Vita-Finzi (1966) with permission from The Prehistoric Society. Fig 11.2(b): from The Mediterranean Valleys: Geological Changes in Historical Times, p. 92, Vita-Finzi (1969), Cambridge, Cambridge University Press, with permission from the author. Fig 11.4(b): from The Mediterranean Valleys: Geological Changes in Historical Times, p. 10, Vita-Finzi (1969), Cambridge, Cambridge University Press, with permission from the author.

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Fig 11.13: from Journal of Field Archaeology, 17, 379–96, ‘Land use and soil erosion in Prehistoric and Historical Greece’, van Andel et al. (1990). Reproduced from Journal of Field Archaeology with permission of the Trustees of Boston University. All rights reserved. Fig 11.14: from ‘Palaeohydrological changes in the Mediterranean region during the Late Quaternary’, Benito p. 131, in Palaeohydology: Understanding Global Change, Gregory and Benito (eds.) (2003). Reproduced with permission of John Wiley & Sons Limited. Fig. 11.16: from Catena 66, 145–54, ‘Past hydrological events reflected in the Holocene fluvial record of Europe’, Macklin et al. (2006) with permission from Elsevier. Fig 11.20: from Journal of the Geological Society, London 162, 879–908, ‘Recent evolution of a Mediterranean deltaic coastal zone: human impacts on the Inner Thermaikos Gulf’, Kapsimalis et al. (2005) with permission from The Geological Society of London. Table 12.2: from Episodes: Journal of International Geoscience 28, 85–92, ‘A formal stratigraphical approach for Quaternary glacial records in mountain regions’, Hughes et al. (2005) with permission from the International Union of Geological Sciences. Fig. 12.8: from Zeitschrift für Gletscherkunde und Glazialgeologie 31, 199–206, ‘Little Ice Age glacier fluctuations in the Pyrenees’, Grove and Gellatly (1995) with permission from the publishers: Universitätsverlag Wagner, Innsbruck. Fig. 12.9: from Little Ice Ages: Ancient and Modern (Volume 1), p. 190, Grove (2004), London, Routledge with permission from the publishers from Cengage Learning Services Limited. Fig. 12.19: from Global and Planetary Change, 50, p. 94, ‘Late Pleistocene glaciers and climate in the Mediterranean’, Hughes et al. (2006) with permission from Elsevier. Table 13.1: from Futures for the Mediterranean Basin: The Blue Plan, p. 31, Grenon and Batisse (1989) with permission from Oxford University Press. Fig. 13.1(a) and (b): maps designed by Dr Thomas Dewez based on the following datasets for topography (STRM30), bathymetry (ETOPO2), and earthquake locations (NEIC). Reproduced with permission from Dr Thomas Dewez (BRGM), French Geological Survey, Orléans, France. Fig. 13.3: from ‘Littoral cells’, Inman 594-599, in Encyclopedia of Coastal Science Schwartz (ed.) (2005). Reproduced with kind permission of Springer Science and Business Media and based on an original figure in: Proceedings of the 19th Coastal Engineering Conference, American Society of Civil Engineers, 2, 1600–17, ‘The Nile littoral cell and man’s impact on the coastal zone of the southeastern Mediterranean’, Inman and Jenkins (1984). Fig. 13.5: from Quaternary Science Reviews, 24, 1969–88, ‘Sea-level change in the Mediterranean since the LGM: model predictions for tectonically stable areas’, Lambeck and Purcell (2005) with permission from Elsevier. Fig. 13.6: from Puglia 2003—Final Conference Project IGCP 437, Sea level change at Capo Caccia (Sardinia) and Mallorca (Balearic Islands) during oxygen isotope sub-stage 5e, based on Th/U datings of phreatic overgrowths on speleothems’, Tuccimei et al. (2003) and originally published in Geodinamica Acta 15, 113–25, ‘Phreatic overgrowths on speleothems: a useful tool in structural geology in littoral karstic landscapes. The example of eastern Mallorca (Balearic Islands)’, Fornós et al. (2002) with permission from Elsevier. Fig. 13.7: from Marine Geology 167, 105–26, ‘Holocene tectonic uplift patterns in northeastern Sicily: evidence from marine notches in coastal outcrops’, Rust and Kershaw (2000) with permission from Elsevier. Fig. 15.12: Reproduced by kind permission of the Syndics of Cambridge University Library. Figs.18.9(a) and (b): Reproduced by kind permission of the Centre Méditerranéen de l’Environment (http://www.cme-cpie84.org/) Fig. 18.14: Reproduced by kind permission of the Conselleria de Medi Ambient del Govern de les Illes Balears. Figs. 19.4 and 19.8: reproduced from open access data from the European Forest Fire Information System (EFFIS) at the European Institute for Environment and Sustainability, EC Joint Research Centre:

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Preface and Acknowledgements Figs. 20.1, 20.5 and 20.10: from Panarchy: Understanding Transformations in Human and Natural Systems, Gunderson and Holling (eds.) (2002) with permission from Island Press, Washington. Fig. 20.4: from Geoderma, 15, 61–70, ‘Vegetal cover to estimate soil erosion hazard in Rhodesia’, Elwell and Stocking (1976) with permission from Elsevier Science Fig. 20.6(b): from ‘The problem’ Kirkby, 1–16, in Soil Erosion, Kirkby and Morgan (eds.) (1980). Reproduced with permission of John Wiley & Sons Limited. Fig. 20.6(c): from ‘The effect of land use on soil erosion and land degradation under Mediterranean conditions’, Kosmas et al. 57–81, in Mediterranean Desertification: A Mosaic of Processes and Responses, Geeson et al. (eds.). Reproduced with permission of John Wiley & Sons Limited. Fig. 20.9: from ‘Erosion-vegetation competition in a stochastic environment undergoing climatic change’, Thornes and Brandt p. 314, in Environmental Change in Drylands: Biogeographical and Geomorphological Perspectives, Millington and Pye (eds.) (1994). Reproduced with permission of John Wiley & Sons Limited. Fig. 23.6: from Biology and Wildlife of the Mediterranean Region edited by Blondel and Aronson (1999) with permission from Oxford University Press. Full references for these figures and tables have been included in their respective chapters. Credits for photographs are given in the respective captions. While every reasonable effort has been made to trace and contact copyright holders, this has not always been successful. We apologize for any apparent negligence. If this list contains errors or inconsistencies, please contact the editor so that these can be corrected in any future editions.

Contents List of Figures List of Tables List of Contributors

xvii xxx xxxii I. The Physical and Biological Framework

Editorial Introduction

3

JAMIE WOODWARD

1. Tectonic Setting and Landscape Development

5

ANNE MATHER

2. The Marine Environment: Present and Past EELCO ROHLING , RAMADAN ABU - ZIED , JAMES CASFORD , ANGELA HAYES ,

33

AND BABETTE HOOGAKKER

3. The Climate System ANDREW HARDING , JEAN PALUTIKOF , AND TOM HOLT

69

4. Cenozoic Climate and Vegetation Change

89

CHRONIS TZEDAKIS

5. The Nature and Origin of the Vertebrate Fauna

139

JACQUES BLONDEL

II. Process and Change in Specific Environments Editorial Introduction

167

JAMIE WOODWARD

6. Weathering, Soils, and Slope Processes

169

JOHN WAINWRIGHT

7. Vegetation and Ecosystem Dynamics

203

HARRIET ALLEN

8. Hydrology, River Regimes, and Sediment Yield JOHN THORNES , FRANCISCO LÓPEZ - BERMÚDEZ , AND JAMIE WOODWARD

229

9. Lakes, Wetlands, and Holocene Environmental Change

255

NEIL ROBERTS AND JANE REED

10. Karst Geomorphology and Environmental Change

287

JOHN LEWIN AND JAMIE WOODWARD

11. River Systems and Environmental Change

319

MARK MACKLIN AND JAMIE WOODWARD

12. Glacial and Periglacial Environments

353

PHILIP HUGHES AND JAMIE WOODWARD

13. Coastal Geomorphology and Sea-Level Change

385

IAIN STEWART AND CHRISTOPHE MORHANGE

14. Aeolian Processes and Landforms ANDREW GOUDIE

415

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Contents

III. Hazards Editorial Introduction

433

JAMIE WOODWARD

15. Volcanoes

435

CLIVE OPPENHEIMER AND DAVID PYLE

16. Earthquakes

469

STATHIS STIROS

17. Tsunamis

493

GERASSIMOS PAPADOPOULOS

18. Storms and Floods

513

MARÍA DEL CARMEN LLASAT

19. Wildfires

541

FRANCISCO LLORET , JOSEP PIÑOL , AND MARC CASTELLNOU

IV. Environmental Issues in the 21st Century Editorial Introduction

561

JAMIE WOODWARD

20. Land Degradation

563

JOHN THORNES

21. Water Resources

583

JEAN MARGAT

22. Air Pollution and Climate

599

JOS LELIEVELD

23. Biodiversity and Conservation

615

JACQUES BLONDEL AND FRÉDÉRIC MÉDAIL

Index

651

List of Figures 1.1. (a) The Alpine Himalayan orogen in its global setting and (b) the main tectonic landform features of the Mediterranean 1.2. Simplified cross-section of subduction rollback 1.3. The present geodynamic framework of the Mediterranean 1.4. Seismic activity in the Mediterranean 1.5. Volcanic activity in the western Mediterranean over the last 33 Ma 1.6. Eastward migration of the topography in conjunction with the eastward rollback of the subduction zone in the Apennines of Italy through time 1.7. An example of alluvial fans from a faulted mountain front in the Tabernas basin of south-east Spain 1.8. An example of well-developed badlands in Tortonian marls within the Tabernas basin of south-east Spain 1.9. Volcanic activity and tsunami impact in the Aeolian Islands 1.10. Plate boundaries and the main poles of rotation 1.11. Oblique view of the Megara basin 1.12. Distribution of erosional and landslide features in relation to the 70 ka river capture site in the Río de Aguas basin, south-east Spain 1.13. Image of an active mass failure along the margins of the Río de Aguas valley 1.14. Schematic evolution of the valley systems of the Sorbas basin before and after the 70 ka river capture 1.15. Surface lowering above the 70 ka river capture site depicted in Figure 1.12 2.1. Map of the Mediterranean Sea 2.2. Longitudinal cross-section showing water mass circulation in the Mediterranean Sea during the present-day winter 2.3. Surface water circulation in the Mediterranean Sea 2.4. Northern Hemisphere summer atmospheric circulation pattern 2.5. Schematic illustration of surface circulation in the Alboran Sea 2.6. Schematic illustration of the main gyres associated with Atlantic surface flow 2.7. Typical salinity profiles for the western and eastern Mediterranean basins 2.8. Schematic illustration of the preconditioning, violent mixing, and deep convection phases 2.9. Estimated sea surface salinity distribution for (a) the Holocene Climate Optimum and (b) the Last Glacial Maximum 2.10. Annual sea surface temperature (SST) reconstructions for the Holocene Climate Optimum and Last Glacial Maximum 2.11. Scanning electron microscope images of the carbonate shells of several planktonic foraminiferal species that live in the Mediterranean Sea 2.12. Example of a laminated sapropel in a freshly opened sediment core 2.13. Schematic presentation of the changes in subsurface circulation patterns between the present day and times of sapropel formation

6 8 9 12 13

17 18 19 20 21 22 24 25 25 26 34 34 35 38 41 41 42 43 48 49 50 51 52

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List of Figures

2.14. Phase relationships between the sapropel record and associated ‰18 O record from core RC9-181 and the orbital cycles of precession and eccentricity 3.1. The location of the Mediterranean region in relation to the large-scale atmospheric circulation 3.2. Seasonal temperature and rainfall variations at selected sites around the Mediterranean 3.3. (a) Mean annual rainfall and (b) length of the dry season across the Mediterranean basin 3.4. (a) Regions of cyclone genesis and dominant cyclone tracks in the Mediterranean and (b) a TERRA satellite image of a cyclone centred on the Ionian Sea 3.5. Composite graphs of annual, winter, and summer temperature in the whole Mediterranean and the western, central, and eastern basins, 1960–2000 3.6. Composite graphs of annual, winter, and summer precipitation in the whole Mediterranean and the western, central, and eastern basins,1960–2000 3.7. Future changes in temperature and precipitation over the Mediterranean 3.8. Future changes in (a) length of the summer drought and (b) maximum five-day precipitation over the Mediterranean 4.1. Location of sites discussed in Chapter 4 4.2. Global oxygen isotope record based on data from more than forty DSDP and ODP sites 4.3. A section of the compiled oxygen isotope record of Figure 4.2 for the interval 4–7 Ma 4.4. Northern Hemisphere palaeogeography and global vegetation maps for selected time slices in the Tertiary 4.5. Summary pollen diagram showing the main taxa of the interglacial succession during the Last Interglacial at Ioannina 4.6. (a) Variations in ‰18 O composition of benthic foraminifera in V19-30 in the East Pacific. (b) Variations in ‰18 O composition of ice in Greenland Ice Sheet Project 2 record. (c) Variations in alkenone-derived sea surface temperatures in marine core MD95-2043 from the Alboran Sea, western Mediterranean. (d) Interval of maximum lake levels of Lake Lisan, Dead Sea Transform area. (e) Interval of Kastritsa beach deposits, Ioannina basin, north-west Greece. (f) Temperate tree pollen percentages curves from Ioannina 1—284, Kopais k93, central Greece 4.7. June insolation for 65˚ N and variations in ‰18 O composition of benthic foraminifera over the last 3 Myr in the Shackleton 06 (S06) composite record from sites in the equatorial East Pacific 4.8. SPOT imagery of the Ioannina basin and surrounding areas showing the extent of topographical variability in the region 4.9. Location of some Mediterranean pollen records from wetland sites spanning all or part of the LGM, with inferred refugial tree populations 4.10. ‘Serial extinction’ of a number of genera in Europe 4.11. Variations in ‰18 O composition of benthic foraminifera in the S06 composite record and arboreal pollen percentages at Tenaghi Philippon, north-eastern Greece

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4.12. Variations in ‰18 O composition of benthic foraminifera in the S06 composite record over the interval 960–1,340 ka 4.13. Variations in ‰18 O composition of ice in Greenland Ice Sheet Project 2 record, and of planktonic foraminifera in core MD95-2042 from the Portuguese margin; variations in alkenone-derived sea surface temperatures in western Mediterranean; temperate tree pollen percentages in marine core MD95-2043 from the Alboran Sea, Ioannina I-284, Kopais K93, and Tenaghi Philippon TF II, Greece 5.1. The four quadrants of the Mediterranean region 5.2. The Lesser kestrel Falco naumanni, a typical but declining species found in old cities and craggy areas in the Mediterranean 5.3. Phylogeography of the Brown bear 5.4. The Brown bear Ursus arctos 5.5. The Ibex Capra ibex 5.6. Patterns of invasion of the Mediterranean basin and western Europe by the house mouse 5.7. Biogeographical origin of the bird fauna of the Mediterranean region 5.8. The Fan-tailed warbler Cisticola juncidis 5.9. Four species of warbler that are typical of Mediterranean matorrals where they evolved: the Mediterranean warbler Sylvia melanocephala; the Subalpine warbler S. cantillans; the Dartford warbler S. undata; and Marmora’s warbler S. sarda 5.10. The Rock partridge Alectoris graeca 5.11. Relationships between the three main groups of Mediterranean warblers (genus Sylvia) and their geographical range 5.12. The zonation of the various vegetation belts in the western Mediterranean area in relation to both altitude and latitude 5.13. Examples of nest types for terns and waders that can be found on small islets within Mediterranean lagoons between April and July 5.14. The Blue tit Cyanistes caeruleus 5.15. The Mediterranean as a key place for migratory and wintering birds 5.16. The Short-toed eagle Circaetus gallicus 5.17. Pygmy hippos and elephants 5.18. Turnover of non-volant mammal species in Corsica as a result of human colonization around 9,500 years ago 5.19. The very large and beautiful Eyed lizard Lacerta lepida 5.20. Levels of endemism of freshwater fish of the large peninsulas of the northern part of the Mediterranean basin 6.1. Potential climatic controls on weathering processes 6.2. Gorges of the River Hérault in south-west France 6.3. Comparison of soil depths as measured in different climate zones and on different lithologies on the island of Lesvos 6.4. Processes leading to the formation of iron oxides and thus the development of brown (goethite) and red (haematite) soils 6.5. Spatial distribution of soils in the Mediterranean basin according to the FAO classification scheme 6.6. Examples of typical soil profiles from the Mediterranean 6.7. Examples of splash pillars forming on marls (the ‘Terres Noires’) near Propiac, south-east France 6.8. Examples of the development of patchy vegetation representing ‘islands of fertility’, from the Montpellier Garrigue, southern France

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6.9. Measured soil displacements as a result of tillage erosion using a mouldboard plough in Tuscany 6.10. Accelerated erosion related to slope-parallel tillage in vineyards in southern France, following the storms of 12–13 November 1999 6.11. Soil production by isopods and porcupines on hillslopes at Sede Boquer, southern Israel 6.12. Comparison of relative interrill erosion rates as a function of amount and type of rock-fragment cover of the soil surface 6.13. Aerial photograph of Stipa tenacissima vegetated slopes at the Rambla Honda, southern Spain 6.14. Rills formed during the extreme rainfalls of 22 September 1992 in south-eastern France 6.15. Badlands at Tabernas, Almería, Spain 6.16. Map of observed pipes showing the extent of subsurface erosion in the Agri basin, southern Italy 6.17. Linked pipe and gully erosion, Murcía, Spain 6.18. Map of erosion and deposition following an extreme storm event of 215 mm in Catalonia, Spain in June 2000 6.19. Examples of different types of mass movement 6.20. Landslide inventories for areas in central Italy 6.21. Large rock slides in the valley of the River Guadalfeo, Granada, Spain 6.22. Idealized evolution of ground-surface properties and surface instability 6.23. Analysis of landslide events that resulted in fatalities in Italy 6.24. Details of the Vaiont landslide disaster 7.1. Mediterranean vegetation communities and trajectories of change 7.2. Bioclimatic life zones of the Mediterranean region 7.3. Invasion of sclerophyllous maquis vegetation into an old olive orchard, Crete 7.4. Typical garrigue vegetation in Crete 7.5. Extent of maquis communities across the Mediterranean region 7.6. Distribution maps for circum-Mediterranean taxa, Olea europea subsp. oleaster, Arbutus unedo, Cistus salvifolius, Lavandula stoechas, and vicariant taxa, Quercus ilex and Q. calliprinos, Pinus halepensis, P. brutia, Quercus suber, and Cercis siliquastrum 7.7. Cistus ladanifer-dominated scrub vegetation of the Algarve, Portugal 7.8. Theoretical degradation and regeneration sequences for primary maquis or sclerophyllous evergreen shrub communities 7.9. The typical ‘hedgehog’ shape of the alpine, Euphorbia acanthothamnos, growing among rocky scree of the Psilorites Mountains, Crete 7.10. Olive terraces on Crete 7.11. Basal regrowth of Arbutus unedo in the spring of 2004, following fire in the summer of 2003, Monchique, southern Portugal 7.12. A goat browsing on Quercus coccifera, Psilorites, Crete 7.13. Spread of evergreen sclerophyllous vegetation across the Mediterranean region during the Holocene 8.1. An upland river catchment in the mountains of north-west Greece 8.2. The water balance of the Mediterranean region showing the major fluxes between the main components of the hydrological cycle 8.3. Total annual runoff from river basins in each country bordering the Mediterranean Sea

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8.4. Eagleson’s decomposition of the annual water balance for catchments in different climatic settings 8.5. Rainfall regimes in the Mediterranean region 8.6. The seasonally dry gravel-bed channel of the Voidomatis River upstream of the Vikos Gorge in north-west Greece 8.7. Average monthly flows for rivers around the Mediterranean basin 8.8. The relationship between relative runoff and vegetation canopy cover 8.9. The species used by Garcia-Ortiz in her study of rainfall partitioning by different Mediterranean plants 8.10. Flood flows and erosion in Mediterranean catchments 8.11. Stone-walled terraces on hillslopes near Campanet in central Majorca 8.12. A water cistern directly under the former Greek agora in Ptolomais, Libya 8.13. (a) A newly built check dam in the Sugura River basin, south-east Spain, (b) a sediment-filled reservoir, Valdeinfierno, Murcia, Spain, and (c) oblique air photograph of the town of Puerto Lumbreras and the Rambla Nogalte in the aftermath of the large flood in October 1973 8.14. Map of the Segura River basin (18,800 km2 ), and the impact of reservoir construction on the monthly distribution of flows in the Segura River, Murcia, south-east Spain 8.15. The Río Aguas at Urra in the Sorbas basin, Almeria, south-east Spain, in flood and dry 8.16. Suspended sediment yield from river basins in mountain environments in different climate and vegetation zones 8.17. (a) Gully erosion in soft sediments in central Israel, and (b) a veneer of fresh suspended sediments deposited within the channel zone of the Torcicoda River, central Sicily 8.18. Suspended sediment yield from Moroccan river catchments formed in different rock types 9.1. Maps showing: (a) Mediterranean type climates; (b) the location of the largest lakes in the circum-Mediterranean region; (c) exemplar lake types; (d) location of selected key Holocene palaeolimnological sites 9.2. Lake Pamvotis, Ioannina basin, north-west Greece: an example of a freshwater lake in a karstic intermontane landscape 9.3. Meke Tuzlası , a hypersaline lake occupying a Late Pleistocene crater on the Anatolian plateau, with a new volcanic cone rising through the middle 9.4. Skadar, the largest freshwater lake in the Balkans 9.5. Lago di Pergusa, a small shallow circular lake on Sicily 9.6. Ternary diagram showing the major chemical anion composition of fifty-seven inland lakes in Spain 9.7. The Dead Sea 9.8. Comparative oxygen-isotope curves for three East Mediterranean lake records 9.9. Key selected palaeolimnological indicators and inferred record of Holocene lake-level change in the Laguna de Medina, Cádiz, south-west Spain 9.10. Stratigraphic changes in major pollen types from Birket Ram showing mid–late Holocene cultural and vegetation change in the Golan Heights, along with catchment erosion indicated by magnetic susceptibility.

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9.11. Lakes with contrasting catchment land uses and aquatic ecology in the Middle Atlas region of Morocco: (a) Tigalmamine and (b) Sidi Ali 9.12. Twentieth-century catchment land use and diatom algae species composition from the recent sediments of five Middle Atlas lakes 9.13. (a) Numbers and areas of Mediterranean–Black Sea Ramsar sites by country, and (b) location of selected coastal wetlands and other threatened sites 9.14. The Huleh basin, Israel, showing wetland extent before and after mid-twentieth-century drainage reclamation 9.15. Burdur, a saline lake occupying a tectonic basin in south-west Turkey 9.16. Late twentieth-century urban expansion of Istanbul into the catchments of the Çekmece coastal lagoons 10.1. The distribution of the major outcrops of carbonate rocks (limestones, dolomites, and marble) in the Mediterranean region 10.2. Steep limestone slopes in Kotor Bay on the coast of Montenegro 10.3. Bedding planes and joints in Mediterranean limestones 10.4. Large-scale solution channels in limestone in a formerly glaciated valley, Durmitor Massif, Montenegro 10.5. Geography students from the University of Manchester exploring the cave systems in the gypsum of the Sorbas basin in south-east Spain 10.6. The tectonic setting for the deposition and deformation of Mediterranean limestones 10.7. The development of vadose and phreatic cave systems in a karst drainage system 10.8. A typical karst system in the Mediterranean region showing material inputs, stores, and outputs and associated processes in the vadose and phreatic zones 10.9. Features produced by the precipitation of calcium carbonate in karst environments in Majorca, Spain 10.10. Fine-grained sediment outputs from a Mediterranean karst system 10.11. The vertical zonation of karst landscapes and processes in ‘European folded mountain regions’ 10.12. Two forms of intentional human modification to hillslopes in karst environments in the Dinaric region 10.13. (a) The distribution of poljes in the Dinaric karst region. (b) The formation of three types of poljes under varying structural and hydrogeological conditions 10.14. Two limestone gorges at different stages of development in the Mediterranean region 10.15. Relict karst pinnacles in the White Desert of Egypt 10.16. Three landscape features in the Mediterranean produced by the precipitation of secondary carbonates 10.17. Karst terrain in north-east Majorca showing bare limestone slopes and thick terra rossa soils on the valley floor 10.18. Rockshelter and cave entrance environments can form important sediment sinks and represent a major archaeological resource 10.19. A high resolution oxygen isotope record (‰18 O) from speleothems in Soreq Cave, Israel, spanning the last 140,000 years 10.20. Human activities in the Mediterranean region and their potential impact on non-karst and karst terrains

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11.1. A building under excavation at the archaeological site of Olympia in the valley of the Alfios River in the western Peleponnese, Greece 11.2. (a) The Quaternary sediments of the Louros Valley in Epirus, north-west Greece. (b) The temporal record of channel and floodplain deposition and incision 11.3. Large boulders on the bed of the channel of the Voidomatis River in the Vikos Gorge of north-west Greece 11.4. (a) View looking across the coastal plain in Cyrenaica (north-east Libya) showing the lower course of the Wadi Zewana. (b) Block diagrams showing the evolution of the Late Pleistocene and Holocene alluvial stratigraphy at a trunk stream tributary confluence in wadi systems in Libya 11.5. (a) The lower course of the Wadi Zewana showing a c.25 m thick exposure in Late Pleistocene alluvium. (b) A section showing the Late Pleistocene river sediments with coarse-grained angular gravels exposed at the base 11.6. Dated alluvial units in river systems across the Mediterranean region between c.130 and 10 ka shown in association with two proxy climate records 11.7. Dated alluvial units in river systems across the Mediterranean region between c.65 and 10 ka 11.8. The Pleistocene and Holocene fluvial stratigraphy in the middle and lower reaches of the Voidomatis River basin 11.9. The sediments exposed during the excavations at Boila rockshelter in the lower reaches of the Voidomatis River in north-west Greece 11.10. The deeply incised valley floor and Quaternary terraces of the Río Aguas in south-east Spain 11.11. The Holocene alluvial sediments and terraces in the middle reaches of the Torcicoda River in central Sicily 11.12. The prosperity and depression model of slope stability and soil erosion 11.13. Patterns of Holocene alluviation in Greece for the last 8,000 years 11.14. Patterns of fluvial aggradation and flooding in five Mediterranean countries for the Holocene period 11.15. (a) A summed probability plot for radiocarbon dates from Holocene alluvial records in Spain (11 ka to present). (b) Probability plots for each of the three alluvial depositional contexts shown. (c) Summed probability plots based on the change and mid-point alluvial data sets and the bracketed slackwater flood deposits. (d) A summed probability plot for the radiocarbon dates from the slackwater sediments shown in relation to the North Atlantic drift ice index 11.16. Probability difference curves of radiocarbon dates associated with major flooding episodes in Great Britain, Spain, and Poland 11.17. The flood record for the Aradena Gorge in south-west Crete between 1840 and 2000 based on lichen dating of coarse-grained flood deposits 11.18. Flood histories from five parts of the Mediterranean since AD 1500 11.19. The deeply incised channel zone of the Alfios River in western Greece 11.20. Summary of human impacts on the river channel systems in the lower reaches and delta complex of the Axios, Aliakmon, and Gallikos rivers in north-east Greece over the last century

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12.1. Map of the Mediterranean showing the main mountain areas referred to in Chapter 12 12.2. (a) Snow accumulation to a depth of c.3 m between the villages of Zabljak and Crna Gora in Montenegro. (b) Looking west from the summit of Mount Orjen towards the summit of Subra, Montenegro 12.3. Map of the Zeleni Sneg glacier on Mount Triglav, Slovenia, depicting ice retreat since the mid-nineteenth century 12.4. The Debeli Namet glacier on the northern slopes of Sljeme, Montenegro 12.5. The effects of avalanching on beech trees on Mount Tymphi, Greece 12.6. The distribution of currently glacierized areas in the Pyrenees 12.7. Unglaciated periglacial surface covered in felsenmeer, Ouanoukrim, Atlas Mountains 12.8. Map of the Glacier d’Ossoue, the largest modern glacier in the Pyrenees 12.9. Map of the glaciers of the Maladeta massif where some of the largest modern glaciers in the Pyrenees are found 12.10. Map of the former extent of Pleistocene glacial and nivation features in the Mediterranean 12.11. Moraines at c.1,700 m a.s.l. in the Vourtapa valley above the village of Skamnelli on Mount Tymphi, Greece 12.12. Limestone pavement on Mount Tymphi, Greece 12.13. Cemented till on Mount Tymphi, Greece 12.14. Well-developed screes within the limits of Vlasian Stage glaciers on the southern slopes of Mount Tymphi, Greece 12.15. The extent of Middle and Late Pleistocene glaciers on Mount Tymphi, Greece 12.16. Glacial geomorphological maps of Mount Olympus, north-eastern Greece 12.17. Moraines at c.1,000 m a.s.l. in Duboki Do, above the village of Ubli on Mount Orjen, Montenegro 12.18. Glacial arête between the peaks of Sedlena Greda and Ranisava in the Durmitor mountain area, Montenegro 12.19. Summary pollen percentage curves from the Ioannina I-284 sequence in north-west Greece, spanning the Last Glacial cycle 13.1. Major tectonic structures and associated seismicity of the Mediterranean region 13.2. Coastal morphodynamics of the Mediterranean basins showing the general near-surface water circulation pattern and the locations and attributes of the four major delta shelves 13.3. The Nile littoral cell extends along the south-eastern Mediterranean coast from Alexandria, Egypt, to the Akhziv Submarine Canyon, Israel 13.4. (a) Elevation of the Last Interglacial shoreline. (b) Rates of late Holocene crustal movement. (c) Predictions of global isostatic adjustment made for Mediterranean tide-gauge stations 13.5. Predicted relative sea levels and shorelines across the Mediterranean region at four epochs 13.6. Schematic representation of a littoral karst cave 13.7. (a) Effects of variations of coastal-wave energy on marine-notch formation. (b) A marine notch at Capo Milazzo in north-eastern Sicily 13.8. The absence of a Holocene sea level above present datum is supported by the evidence of painted horses on a wall of a half-submerged Palaeolithic cave near Marseilles 13.9. A schematic coastal profile showing the main characteristics of bioconstruction and biodestruction on calcareous coasts in the western and eastern Mediterranean

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13.10. Age-depth diagram from Marseilles’s archaeological excavations compared with dated algal rims from nearby rocky cliffs 13.11. Measured relative sea-level changes in the old harbour of Pozzuoli compared to estimated relative sea-level changes using biological indicators 13.12. Recent relative sea-level variations in Antikythira island, Greece 13.13. Tsunami activity in the Mediterrenean Sea 13.14. Franchthi Cave in the south-east Peloponnese, Greece 13.15. Historical records of coastal flooding for the Rivers Tiber and Rhône 14.1. A map of some aeolian phenomena and locations in the Mediterranean basin 14.2. The passage of dust systems from North Africa to the Middle East, mid-March 1998 14.3. The TOMS sequence across North Africa to the Middle East for mid-April 2000 14.4. The loess of Matmata, southern Tunisia, has been inhabited by cave dwellers 14.5. Barchan dunes in the Libyan Desert, Kharga depression, Egypt 14.6. Gypsum crust soils with polygonal structures in southern Tunisia 14.7. The great lunette dune on the lee side of the Sebkha el Kelbia, central Tunisia 15.1. Map of the Mediterranean basin showing the locations of selected volcanoes and volcanic provinces 15.2. Volcanic hazards: Mt. Etna erupting in 2001 15.3. Fumarolic and diffuse soil emissions on Vulcano (Italy) pose a health hazard 15.4. Map to show the extent of fallout from the Y5/Campanian Ignimbrite eruption 15.5. View of Herculaneum and modern Ercolano 15.6. Plaster cast of one of Pompeii’s victims 15.7. Sequence of laminated deposits from pyroclastic density currents in the Monte Guardia area of Lipari 15.8. Stromboli volcano 15.9. Volcano seen from Lipari 15.10. Aerial view of Mt. Etna rising above the city of Catania 15.11. The town of Fira clinging to the rim of Santorini’s caldera 15.12. Map of part of Santorini in c.1715, showing the Kameni islands after the 1707 eruption 15.13. (a) Map of Nea Kameni and Mikra Kameni, after the 1866–70 eruptions, and (b) hillshade digital elevation model of the present-day status of the Kameni islands 15.14. (a) The duration of eruptions on the Kameni islands since AD 1570, and (b) the heights of lava domes as a function of time elapsed during an eruption for the 1866 and 1939 Kameni eruptions compared to the domes of Mt. St Helens and St Vincent 15.15. Volcano monitoring and crisis response 16.1. (a) Ruins of buildings in Verneuges, Provence, that were destroyed by the 1909 Lambesc earthquake. (b) A tombstone from the ancient Greek town of Nikomedia, commemorating the death of two young boys and their teacher in AD 120

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16.2. Epicentres of shallow earthquakes in the Mediterranean, 1961–83 16.3. Epicentres of earthquakes, 1900–65, and areas affected by historical earthquakes in the eastern Mediterranean 16.4. Plate boundaries in the eastern Mediterranean 16.5. Seismicity across the Peloponnese 16.6. Examples of earthquake damage to ancient buildings 16.7. Faulting and river response during the 1980 earthquake in southern Italy 16.8. Rocks uplifted during the 1953 earthquakes, Ionian Islands 16.9. Progressive rupturing of the North Anatolian Fault, 1939–99 16.10. Areas affected by the 1202, 1926, and 1927 earthquakes, eastern Mediterranean 16.11. (a) A normal fault produced by the 1954 earthquake in Thessaly. (b) Reverse faulting, folding, and uplift that dammed the Cheliff River during the 1980 Al Asnam earthquake, Algeria. (c) Railway tracks offset by strike-slip faulting during the 1999 Izmit earthquake in Turkey 16.12. Contours of uplift across western Crete resulting from the AD 365 event 16.13. Response of a stream at Sougia, Crete, to an uplift during the AD 365 earthquake 16.14. The effects of ground sliding, compaction, and perhaps liquefaction following the 1783 earthquake in Calabria, Italy 16.15. (a) A collapsed multi-storey building in Kalamata, Greece, following the 1986 earthquake. (b) During the same event many traditional buildings were badly damaged but did not collapse 17.1. Co-seismic dip-slip motion along faults and tsunamigenesis near deep sea trenches 17.2. A schematic explanation of some of the tsunami terms used in Chapter 17 17.3. (a) Tsunamigenic zones in the Mediterranean Sea. (b) Types of field evidence for the occurrence of past tsunamis 17.4. Important tsunamis reported for southern Italy 17.5. Some of the damage caused by the Stromboli tsunami of 30 December 2002 17.6. Important tsunamis along the Hellenic arc 17.7. An excavated section showing palaeotsunami deposits in Dalaman, south-west Turkey 17.8. Important tsunamis in the Cyclades Islands in the southern Aegean Sea 17.9. Palaeotsunami investigation within an archaeological excavation in St George, in eastern Thera 17.10. Detail of the tsunami deposits exposed and attributed to the 30 September 1650 volcanigenic tsunami 17.11. (a) Important tsunamis in the Gulf of Corinth and the Maliakos Gulf. (b) The coast prior to the large landslide that produced a tsunami in the Gulf of Corinth, 1963 17.12. Important tsunamis in (a) the east and north Aegean Sea, (b) the Sea of Marmara, and (c) the Cyprus-Levantine Sea area 17.13. The frequency of tsunamis in the Mediterranean Sea as a function of longitude and latitude, 1628 BC to AD 2003 17.14. The cumulative frequency and intensity of Mediterranean tsunamis as a function of time, 1628 BC to AD 2003 17.15. Relationship between between intensity and frequency of tsunamis for Greece, Italy, and the Mediterranean Sea

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17.16. Tsunami intensity as a function of earthquake intensity and magnitude for the entire Mediterranean Sea 18.1. (a) The key study areas of adverse weather phenomena in this chapter. (b) The key sites and river systems in north-east Spain and southern France cited in the text 18.2. The mean annual frequency of cyclones in the Mediterranean, summer and winter 18.3. Classification of cloud systems associated with heavy rainfall events in the Mediterranean area 18.4. An example of the synoptic situation that produces heavy rainfall with catastrophic floods in the western Mediterranean region 18.5. Rainfall distribution for a flood event of Type 2b, 6–8 November 1982 18.6. Heavy rainfall that produced a flash flood event of Type 1, 3 April 1989 in the Pyrenees 18.7. A flood event of Type 2a, 10 June 2000, in Catalonia 18.8. Destruction of a road bridge by the Riera de la Magarola during the flood event of 10 June 2000 18.9. Some consequences of the flash flooding produced in the Gard region of southern France, 8 September 2002 18.10. A flood event of Type 3, 21–30 January 1996, in Catalonia 18.11. The relationship between maximum flow and average annual catchment flow for river catchments in the Mediterranean and non-Mediterranean regions of Europe 18.12. The record of catastrophic floods in Catalonia since the early fourteenth century 18.13. A ‘Levante’ wind storm that affected Catalonia, 16–18 October 2002 18.14. The destruction of forests in the Balearic Islands by the western Mediterranean ‘superstorm’ of November 2001 18.15. A tornado recorded in Barcelona, on 8 September 2005 19.1. Fires in Mediterranean basin countries 19.2. Temporal variation in (a) the number of fires per year and (b) the area burnt 19.3. The proportion of fires and of area burnt in relation to fire size, France 19.4. The area burnt by forest fires in Portugal during 2003 19.5. A fire scar at the base of a pine tree trunk in Catalonia, north-east Spain 19.6. A conceptual model of factors influencing the area burnt in a region 19.7. (a) Changes in summer climatic fire risk in Catalonia, and (b) the burnt area in relation to the number of days with high fire risk 19.8. Fire risk in Europe according to the Canadian FWI 19.9. Burned forest in a highly populated area of Catalonia 19.10. A fire prevention sign in the uplands of Majorca 19.11. An illustration of a fire regime model 20.1. Different phases of the Holling and Gunderson adaptive cycle 20.2. An application of the USLE for the Autonomous Region of Andalucia 20.3. The EU erosion estimate for Spain 20.4. Soil loss and runoff as a function of the proportion of ground covered by a vegetation canopy 20.5. A metaphor for a system’s stable and unstable conditions 20.6. (a) The relationship between sediment yield and annual effective rainfall. (b) Estimated rates of erosion by wind and water. (c) Erosion for field plots around the Mediterranean

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20.7. The evolution of plant cover over time on abandoned plots in the Spanish Pyrenees 20.8. Runoff and sediment yield data for different land uses on regenerated abandoned fields in the Spanish Pyrenees 20.9. Simulated response of vegetation to stochastic rainfall 20.10. A nested set of adaptive cycles in time and space 21.1. Distribution of potential annual runoff (effective precipitation) in the Mediterranean basin 21.2. The Mediterranean drainage network and river basins 21.3. Natural renewable and exploitable water resources per country in the Mediterranean basin 21.4. Natural renewable water resources and real exploitable water resources per inhabitant in the Mediterranean basin 21.5. The exploitation index of natural renewable water resources across the Mediterranean basin 22.1. European air pollution emissions, 2000 22.2. Widespread aerosol haze in the Mediterranean basin 22.3. Schematic of air flows during MINOS, 2001 22.4. Transport-time spectrum showing the period between release of air pollutants and arrival in the upper troposphere 22.5. Model-calculated ozone at the surface for the present and the possible future 22.6. Mean diurnal cycles of OH and NO3 radicals 22.7. Mean particle composition during MINOS for fine and coarse mode aerosols 22.8. Diurnal mean, clear-sky radiative forcing during MINOS 22.9. (a) Estimated historical SO2 emissions in Europe; (b) 5-year running mean of Mediterranean SST anomalies 22.10. Percentage changes in annual precipitation comparing low and high SST 23.1. The ten regional hotspots of the Mediterranean basin for plant endemism and richness 23.2. A montado in Portugal with cattle and charcoal burners 23.3. Forest recovery on ancient terraces 23.4. The last individual of a former forest of Juniperus thurifera in southern Morocco 23.5. Wetlands, among the most threatened habitats in the Mediterranean region 23.6. The crustacean branchiopod Triops cancriformis with the rare thrumwort Damasonium stellatum and the parsley frog Pelodytes punctatus, Camargue 23.7. A rare amphibian of temporary ponds of the western part of the basin, the Mediterranean newt, Triturus marmoratus 23.8. Urbanization of coastal areas threatens habitats and rare plants and animals 23.9. (a) Mean laying date of the Blue tit Cyanistes caeruleus in mainland and Corsican habitats (b) Variation of clutch size in the Corsican and mainland habitats 23.10. Intense habitat degradation due to repeated fire events results in very low scrubby vegetation and bare ground 23.11. Carpobrotus acinaciformis, a very aggressive invasive plant species in coastal areas

573 573 574 578 584 585 588 589 593 600 601 602 604 605 607 608 609 609 610 616 617 619 620 622

623 624 625

626 627 630

List of Figures

23.12. Ramet of Zelkova sicula (Ulmaceae), a relict and very threatened palaeoendemic small tree, south-eastern Sicily 23.13. Acis fabrei, a narrow endemic with only four known populations from the southern slopes of the Mont Ventoux 23.14. Three taxa of salmonids: Marble trout, Corsica trout, and Brown trout 23.15. The Pond terrapin Emys orbicularis 23.16. A pair of Bonellis’eagles Hierraaetus fasciatus at their nest 23.17. The Scops owl Otus scops, threatened by the decline of large invertebrates 23.18. The Vikos Gorge with its spectacular limestone cliffs 23.19. Cliffs are important habitats for several rare birds and endemic plant species throughout the Mediterranean basin 23.20. Distribution maps within the Mediterranean bioclimatic region of fifty glacial refugia

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632 633 635 636 637 639 641 644 645

List of Tables 1.1. A summary of the impact of tectonics on the geomorphology of the Mediterranean region 1.2. Examples of badlands and partial badlands found across the Mediterranean 2.1. Contributions to total Nile discharge from the main tributaries 3.1. Centres of cyclogenesis in the Mediterranean 3.2. Local winds of the Mediterranean 3.3. Decadal distribution of heatwave days 3.4. Summary of some recent flash floods in Mediterranean Europe 4.1. Scales of environmental variability and vegetation responses 4.2. List of genera of species’ first fossil appearance in the Tertiary record 4.3. A sample of Mediterranean pollen records 5.1. Endemism rates of several groups of species in the Mediterranean 6.1. Comparison of porportions of different soil types according to the FAO classification relative to location in the north or south of the Mediterranean basin 7.1. Terms used to describe Mediterranean sclerophyllous shrubland 7.2. Fire-adapted strategies of some selected Mediterranean taxa 7.3. Flammability of selected Mediterranean plants based on laboratory tests of leaf ignition of Cretan species 8.1. Water and sediment fluxes from the Axios and Aliakmon rivers that drain into the north-west Aegean Sea 9.1. Individual characteristics of all permanent natural lakes in the circum-Mediterranean region >200 km2 in area, excluding coastal lagoons 9.2. Exemplars of Mediterranean lake types 9.3. Selected key Holocene palaeolimnological records for the Mediterranean 9.4. Selected key Mediterranean wetlands requiring conservation or restoration 10.1. Large discharge springs of the world with flows >20 m3 s−1 10.2. The twenty deepest caves in the world 10.3. Characteristics of active karst settings and passive karst settings for rockshelter and cave entrance environments in limestone terrains 10.4. Karst sites in the Mediterranean region with World Heritage status 12.1. Modern glaciers in the Mediterranean 12.2. Correlation table showing the relationship between the fragmentary glacial sequence in the Pindus Mountains, Greece, and the continuous lacustrine parasequence in the nearby Ioannina 249 and 284 cores 12.3. Current understanding of the geochronology of glacial deposits in the Mediterranean region 13.1. Coastal environments around the Mediterranean Sea classified into bedrock and accretion coasts 13.2. Amplitude, duration, permanence, and length of coast affected by various types of rapid relative sea-level change in the Aegean 14.1. Dust over the Mediterranean 14.2. Dust deposition amounts across the Mediterranean

7 19 40 75 78 79 86 90 98 109 158

177 206 218 220 246

257 258 259 260 298 301 310 314 354

369 376 386 401 407 421

List of Tables

14.3. Particle size characteristics of dust in various parts of the Mediterranean 15.1. Major ash layers correlated with known volcanic eruptions in the Mediterranean over the past 200 kyr 15.2. Summary of volcanic hazards 15.3. Twentieth-century record of deaths, injuries, and other impacts of volcanic activity in the Mediterranean 15.4. Fatal eruptions of Somma-Vesuvio 16.1. Some major or otherwise noteworthy earthquakes in the Mediterranean 16.2. A list of indicative criteria for the identification of earthquakes from archaeological data 17.1. Strong tsunamis of intensity k ≥ 4 reported for the Mediterranean Sea between 426 BC and AD 2002 17.2. Mean return period of tsunami intensity and the most likely maximum tsunami intensity to be observed for various parts of the Mediterranean Sea 17.3. Tsunami potential in each of the tsunamigenic zones of the Mediterranean 18.1. Major flood events in the European Mediterranean since 1990 18.2. The number of catastrophic floods based on historical sources recorded in various river basins in Spain, Italy, and France 20.1. National soil erosion risk data for five Mediterranean countries in the EU 21.1. Water resources in the Mediterranean basin by country and continent 21.2. Key figures on internal and external natural and exploitable water resources for the three main regions of the Mediterranean basin 21.3. Annual water withdrawal volumes for the three main regions of the Mediterranean basin 21.4. Present-day pressures on water resources in the Mediterranean basin 21.5. Water demand predictions for 2025 in the three regions of the Mediterranean basin 22.1. Air pollution emissions in Europe in the year 2000, and emission reductions between 1980 and 2000 23.1. Some of the most invasive alien plants occurring in the Mediterranean basin 23.2. Threatened vascular plants by country based on the former IUCN categories and included in the 1997 IUCN Red List of Threatened Plants 23.3. Threatened vascular plants on the seven large Mediterranean islands, based on the former IUCN categories 23.4. List of large mammals that were present in the Mediterranean basin during the Late Pleistocene, including species found as fossils in various deposits of southern France, and that became extinct 23.5. Major protected areas such as National Parks and Biosphere Reserves within the Mediterranean bioclimatic region 23.6. List of Ramsar sites within the Mediterranean bioclimatic region 23.7. Impacts of the major influences on the biodiversity of ten Mediterranean regional hotspots

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421 436 437 441 444 470 475 495

507 508 514 530 566 587 588 593 594 595 600 629 631 631

634 640 640 644

List of Contributors Ramadan Abu-Zied is an assistant lecturer in the Geology Department at Mansoura University in Egypt. He obtained his Ph.D. in 2001 with Eelco Rohling in Southampton. He is a specialist in benthic foraminifera and the application of their abundance variations and stable isotope ratios in palaeoenvironmental research, with emphasis on the eastern Mediterranean. Harriet Allen is a lecturer and researcher in the Department of Geography, University of Cambridge. Her research interests focus on the response of ecosystems to environmental change. This includes the integration of high resolution remote sensing data with ecological surveys to assess contemporary ecosystem changes, and pollen and sedimentological research to reconstruct longer-term ecosystem changes. Much of her recent fieldwork has been carried out in Mediterranean-climate regions, including Greece and Portugal. Jacques Blondel is with the Centre d’Écologie Fonctionnelle et Evolutive, CNRS, Montpellier and he previously taught at the University of Louvain in Belgium. His main research interests focus on the origin and regulation of biological diversity at several scales of space and time, from processes involved in the establishment of faunas at the scale of the Mediterranean region, to community dynamics at the scale of landscapes and populations at the scale of local habitats. He conducts a long-term (>30 years) programme on the phenotypic variation of birds in Mediterranean habitat mosaics. He is also concerned with biodiversity and conservation issues. Maria del Carmen Llasat is a professor and coordinates a research group in the Department of Astronomy and Meteorology at the University of Barcelona. She was president of the Natural Hazards Section in the former European Geophysical Society (now part of the European Geosciences Union). She is the managing editor of the journal Natural Hazards and Earth System Science. Her research interests are mainly meteorological and hydrometeorological risks in the Mediterranean region. James Casford in a lecturer in the Department of Geography at Durham University. He obtained his Ph.D. in 2001 with Eelco Rohling in Southampton. His research focuses on climate variability and the palaeoceanography of marginal basins, particularly the eastern Mediterranean. Marc Castellnou is the Fire Analysis Officer in the Fire Service of the Catalonian Regional Government where he runs the forest fire training programme for forest fire professionals. He has worked and tackled fires in the USA, Africa, France, the UK, and Portugal. He has also carried out research into forest fire regimes, forest fire ecology, and forest fire propagation. He has been involved in several European research projects as researcher and coordinator. More recently he has coordinated the operational training of international fire-fighting units. Andrew Goudie is a Professor of Geography in the University of Oxford and Master of St Cross College. His research interests are in desert geomorphology and climate change. He has worked extensively in Africa, India, and the Middle East on such themes as dunes, dust, pans, loess, and salt. Andrew Harding is a research associate at the Global Environmental and Climate Change Centre (GEC3 ) at McGill University working with Environment Canada. His main focus is the investigation of linear and non-linear links between synoptic scale atmospheric dynamics and meso-scale climate extremes. His Ph.D. (from the Climatic Research Unit

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at the University of East Anglia) focused on patterns and trends evident within the variability and sensitivity of Mediterranean climate extremes. Angela Hayes is with the Department of Geography, Mary Immaculate College at the University of Limerick. She obtained her Ph.D. in 1999 with Eelco Rohling in Southampton. She specializes in reconstructing past sea surface temperatures from planktonic foraminiferal abundance data, and uses this in combination with stable isotope ratios to reconstruct past ocean and climate conditions. Tom Holt is a Senior Research Associate in the Climatic Research Unit, University of East Anglia. His main research interests are the assessment of uncertainty in projections of climate change from numerical models and the analysis of climate extremes. A special focus has been on likely changes in climate extremes over the Mediterranean to 2100, with a particular emphasis on drought and impacts on the tourism, agriculture, and energy industries. Babette Hoogakker is a postdoctoral researcher in the Department of Earth Sciences at the University of Cambridge. She obtained her Ph.D. in 2003 with Eelco Rohling in Southampton. She is specialized in the combined use of sedimentology, foraminiferal census counts, and shell chemistry, to reconstruct palaeoceanographic conditions on global scales. Philip Hughes is a lecturer in Physical Geography at the University of Manchester, where he held a postdoctoral fellowship between 2004 and 2006 working on the glacial and periglacial history of the Mediterranean. His Ph.D. (University of Cambridge, 2004) focused on Quaternary glaciation in the Pindus Mountains of north-west Greece and he has also worked in Montenegro and Morocco. Jos Lelieveld is director of the Max Planck Institute for Chemistry, and is professor in Atmospheric Physics at Mainz University. His research addresses photo-oxidants (e.g. ozone), the cleaning mechanism of the atmosphere, aerosols, and links with climate. John Lewin is Emeritus Professor of Physical Geography at Aberystwyth University. He is a fluvial geomorphologist who has been concerned with the development of river landforms over a wide range of timescales: long-term landscape evolution dating back to the Tertiary, Quaternary alluvial deposits, historical river channel changes, and contemporary river processes. Research over the more recent timescales has especially involved human impacts and the dispersal of polluted sediments. He has wide field experience of the Mediterranean region. He co-edited Mediterranean Quaternary River Environments (1995) with Mark Macklin and Jamie Woodward. Francisco Lloret is Professor of Ecology at the Universitat Autònoma Barcelona, and researcher at the CREAF (Centre for Ecological Research and Forestry Applications, Spain). His research interests focus on the structure and dynamics of plant communities in relation to anthropogenic sources of disturbance, such as fire regime, land use change, climate change, and exotic plant invasion. He has studied the historical patterns of fire regime in the Mediterranean basin and the interaction between fire regime and vegetation recovery after fire. He has worked in the Mediterranean basin, the United States, Mexico, and Australia. Francisco López-Bermúdez is a Professor in the Department of Physical Geography at the University of Murcia. He published extensively on the hydrology and dynamics of fluvial systems in semi-arid environments. He is especially interested in the interactions between vegetation and erosion in Mediterranean river basins and in the generation and geomorphological impact of large floods. In 2006 (with Jorge García-Gómez) he published Desertification in the Arid and Semiarid Mediterranean Regions. A Food Security Issue in the NATO Security through Science Series.

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Mark Macklin is Professor of Physical Geography at Aberystwyth University. His principal research interest is studying fluvial system responses to short- and longer-term environmental change. His work on Mediterranean rivers has focused on the Iberian Peninsula, mainland and island Greece, and Libya. These investigations have documented the effects of Late Pleistocene glaciation and sub-orbital scale climate change on river behaviour, established the geomorphic impacts of extreme Holocene flood events on mountain catchments and bedrock gorges, and developed a range of remediation and management strategies for river basins contaminated by historical and present-day metal mining. He has just started a three-year project investigating the role of tectonics on historical river development in western Crete. Jean Margat is a geology graduate who worked as a hydrogeologist from 1947 to 1989, first in Morocco (Geological Survey), then in the ‘Bureau de Recherches Geologiques et Minières’ (BRGM: French Geological Survey) focusing on groundwater research, in France and in the arid zone. Later he specialized in water resources in the Mediterranean basin. He is the author of numerous publications dealing with the assessment and the management of water resources, the most recent being ‘Water for the Mediterranean, Present and Future’ (Mediterranean Action Plan /Blue Plan, 2004). He is vice-chairman of the Association ‘Blue Plan for the Mediterranean’ and the ‘Mediterranean Institute of Water’. Anne Mather is a Reader in Earth Sciences with the School of Geography at the University of Plymouth. Her research focuses on long-term landscape development in drylands with the main focus on tectonic geomorphology. This research encompasses both direct and indirect responses of alluvial and fluvial systems to regional tectonics. Her main geographical areas of research include Spain, Turkey, Morocco, and Chile. Frédéric Médail is Professor of Plant Ecology and Biogeography in the Mediterranean Institute of Ecology and Palaeoecology (IMEP, University Paul Cézanne Aix-Marseille III). His research interests include the conservation and biogeography of Mediterranean plants, the processes induced by biological invasion and insular ecology. He conducts his research at several ecological scales, from regional phylogeography to the population biology and ecology of rare and endemic plants better to understand patterns and processes involved in the diversity of the Mediterranean basin hotspot. Christophe Morhange is Professor of Physical Geography at the University of Provence and a member of the CEREGE’s (CNRS) Geomorphology and Tectonics group, Aix-enProvence, France. His research interests are: (1) Holocene relative sea-level changes using biological indicators; and (2) coastal geoarchaeology, notably the use of ancient harbour archives to reconstruct natural and anthropogenically forced changes since antiquity. He has worked in numerous sites around the Mediterranean and the Black Sea, including Bulgaria, Cyprus, Egypt, France, Greece, Israel, Italy, Lebanon, Spain, Tunisia, and the Ukraine. Clive Oppenheimer is based at the Department of Geography, University of Cambridge. His research focus is the development and application of remote sensing techniques for environmental monitoring, especially volcanology. His observations of volcanic gas and aerosol emissions have been used to investigate the transport of magma below volcanoes, as well as the impacts of volcanic pollution on the atmospheric environment. Recently, he has studied the lava lakes of Mt. Erebus in Antarctica, and Erta ’Ale in Ethiopia. He is also interested in the climatic and human impacts of major historic and prehistoric eruptions. Jean Palutikof is Head of the Technical Support Unit, IPCC Working Group II (Impacts, Adaptation, and Vulnerability). She is based in the Hadley Centre at the UK Met Office. Previously, she worked in the Climatic Research Unit at the University of East Anglia, and in the Department of Geography at the University of Nairobi. Her research interests

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focus on climate change impacts, and the application of climatic data to economic and planning issues. She worked on the EU-funded MEDALUS projects, constructing scenarios of regional climate change for the Mediterranean region. Gerassimos Papadopoulos is Research Director with the Institute of Geodynamics, National Observatory of Athens, Greece. His main research interests are in instrumental and historical seismicity, earthquake prediction, and tsunamis, particularly in the EuroMediterranean region. He has worked as a visiting scientist at MIT (Boston, 1984), NIED (Tsukuba, Japan, 1993), and Tohoku University (Japan, 2004. He served as President of the International Natural Hazards Society (2000–6) and he is Vice-President of the European Seismological Commission. Josep Piñol is at the Centre for Ecological Research and Forestry Applications (CREAF) at the Autonomous University of Barcelona. His first work on forest fires focused on the development of methodologies to estimate fire risk, in particular using meteorological indices and measurements of the moisture content of fine fuels. More recently, he has focused his attention on understanding fire regimes in Mediterranean regions, and, in particular, the role of fuel build-up in increasing the occurrence of very large fires. David Pyle is Professor of Earth Sciences in the Department of Earth Sciences at Oxford University and previously taught at the University of Cambridge (1991–2006). His principal interests are understanding patterns and processes of active volcanism, in particular the dispersal of tephra during large eruptions; the emission and reactivity of gases from volcanic vents; and the interactions between volcanoes and the climate system. He has worked in Greece, Italy, Russia, the Americas, and south-east Asia. Jane Reed is a lecturer in the Department of Geography, University of Hull. Her research interests focus on the use of lacustrine diatoms as quantitative environmental indicators in fresh and saline lakes, both in the palaeolimnological study of lake sediment cores and in environmental biomonitoring. Her main geographical focus is the lakes of the Mediterranean and the Balkans, encompassing research themes ranging from long-term Quaternary climate change to recent water pollution and desiccation. Neil Roberts is Professor of Physical Geography at the University of Plymouth. He received his Ph.D. from the University of London (UCL), and has been a researcher at the University of Oxford and subsequently Lecturer at Loughborough University. His research emphasizes global change during the Late Quaternary period, specifically lake sediment-based archives of past climate variability in low and mid-latitude regions, with links to archaeology. He is author of the key text, The Holocene, published by Blackwell. Eelco Rohling is a Professor at the School of Ocean and Earth Science, Southampton University, and is based at the National Oceanography Centre, Southampton. He works on the processes of ocean and climate change over a range of timescales, with emphasis on Pleistocene and Holocene records from subtropical marginal seas, such as the Mediterranean and the Red Sea. Iain Stewart is Professor of Geoscience Communication at the University of Plymouth. A former president of the INQUA Commission on Neotectonics, his primary research interests are on Holocene coastal tectonics and sea-level change, with particular emphasis on shoreline records of earthquake, volcano, and tsunami activity. With Claudio VitaFinzi, he co-edited The Geological Society of London’s Special Publication 146 on Coastal Tectonics, and has studied Holocene coastal change in tectonically active parts of the Mediterranean, principally Aegean Greece and Turkey, and eastern Sicily. Stathis Stiros is Associate Professor with the Department of Civil Engineering, Patras University, and was previously with the Institute of Geology and Mineral Exploration (IGME)

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in Athens. His research interests include the identification of earthquakes from archaeological and coastal geomorphological data and the modelling of crustal and ground deformation processes. He is also interested in new developments in geodetic instrumentation and analytical tools for the recording and analysis of small-scale oscillatory movements, as well as geodetic techniques used in antiquity. John Thornes, who died in 2008, was Research Chair in Physical Geography at King’s College, University of London. His main research was in the interaction of grazing, vegetation and erosion. He has jointly published Environmental Issues in the Mediterranean with John Wainwright for Routledge. He was awarded Doctor Honoris Causa by the University of Murcia in 2006 for his contributions to the knowledge of the environments of southeast Spain. Chronis Tzedakis is Professor of Global Change Palaeoecology at the Earth and Biosphere Institute, School of Geography, University of Leeds. His research centres on understanding the response of vegetation to variations in climatic forcing on different timescales (orbital and millennial/centennial) in the Mediterranean region. His work involves the study of long lake sequences and deep-sea cores, which provide an opportunity to examine phase and amplitude relationships between climate and vegetation changes over several glacial–interglacial cycles. John Wainwright is Professor of Physical Geography at the University of Sheffield, and was previously at King’s College London. His research focuses on the role of erosion processes in land degradation over a range of timescales from prehistory to the modern day, using combined field, laboratory, and computer-modelling approaches. He has worked extensively in southern France, Spain, Italy, and Greece, as well as in dryland environments in the USA and Africa. Jamie Woodward is Professor of Physical Geography at the University of Manchester. His research focuses on Quaternary environmental change, geomorphological systems, and geoarchaeology in the Mediterranean. He has field experience in various parts of Greece, Sicily, Montenegro, Corsica, and Spain. He has also worked in the Nile Valley with the British Museum. He is the co-editor of Geoarchaeology: An International Journal. He is especially interested in fluvial and glacial archives of change and has also worked on Late Pleistocene and Holocene sedimentary records in Mediterranean rockshelters and caves.

I

The Physical and Biological Framework

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Editorial Introduction Jamie Woodward

By examining both contemporary processes and longterm records of change, this volume explores the climates, landscapes, ecosystems, and hazards that comprise the Mediterranean world. This is the only region on Earth where three continents meet and their interaction has produced a very distinctive physical geography. This book examines the landscapes and processes at the margins of the three continents and the distinctive marine environment between them. In broad terms, the physical geography of the Mediterranean is a product of long-term interplay between tectonic forces, climate change, river basin and marine processes, and biosphere dynamics, as well as the action of humans during the course of the Holocene. From the outset, it is important to keep in mind that this physical geography is an integration of energy, materials, and processes within a much wider global system. The Mediterranean is a zone of convergence and interaction. It is a meeting place not only for tectonic plates, but also for air masses, energy, and river flows from both temperate and tropical latitudes. The region also interacts directly with the global ocean, receiving cool North Atlantic waters in exchange for the warmer and saltier waters produced in the basins of the Mediterranean Sea. It is also a biodiversity hotspot; the Mediterranean has been a meeting place for plants, animals, and humans from three continents throughout much of its history. The chapters in Part I set out the physical and biological framework for the rest of the book and examine key debates about the evolution of the Mediterranean environment. They explore fundamental interactions between the lithosphere, atmosphere, hydrosphere, and biosphere across a range of spatial and temporal scales. The scene is set for later chapters that focus more closely on particular aspects of the Mediterranean environment such as ecosystem dynamics, river basin systems, karst environments, natural hazards, and land degradation.

Chapter 1 examines the role of tectonic processes in the development of the Mediterranean landscape and its marine basins. Also highlighted are the dramatic environmental changes and the geomorphological legacy associated with the Messinian Salinity Crisis of the Late Miocene. Chapter 2 focuses on the marine environment, both ancient and modern. The Mediterranean Sea is a relatively small body of water in the global ocean system that has reacted in a sensitive way to environmental change; this is well illustrated by the repeated formation of sapropels throughout the Quaternary. The climate system is the focus of Chapter 3 and this completes the trio of opening chapters that explore the land, sea, and atmosphere of the Mediterranean and some of the key interactions between them. Chapters 4 and 5 examine major themes and ideas in Mediterranean biogeography by exploring the nature and evolution of the region’s vegetation and vertebrate fauna respectively. Chapter 4 combines an analysis of the long-term record of Cenozoic climate and vegetation change with a detailed evaluation of the region’s long Quaternary vegetation records. Together, Chapters 2 and 4 show how key components of the Mediterranean environment responded to the rhythms of global change during the Quaternary. The region provided important refugial areas for plants, animals, and humans during the cold stages of the Quaternary. Pollen records show that trees expanded and contracted their ranges very rapidly in response to abrupt changes in the climate system. Such ecological changes were part of a highly dynamic Quaternary geography that tracked the major environmental fluctuations of the Northern Hemisphere. These themes are developed further for specific environments and settings in the chapters of Part II in conjunction with assessments of the nature and significance of human activity across the Mediterranean in the Holocene.

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1

Tectonic Setting and Landscape Development Anne Mather

Introduction The Mediterranean is the westernmost part of the globalscale Alpine-Himalayan orogenic belt which stretches from Spain to New Zealand (Figure 1.1a). The landscapes of the region have a long and complex history that includes both horizontal and vertical crustal movements and the creation and destruction of oceans (Table 1.1). This began with the break up of the supercontinent Pangea around 250 Ma, which generated the Tethys Ocean—the forerunner to the present-day Mediterranean Sea. Collision of the African and European tectonic plates over the last 30 Ma led to the destruction of the Tethys Ocean, although a few remnants of its geology are preserved within the eastern Mediterranean. It is the collision of Africa and Eurasia, and the associated tectonics that have been largely responsible for generating the Mediterranean Sea, its subsequent history, and the landscapes that surround it. This collisional history progressively reduced the connectivity of the Mediterranean Sea with surrounding marine bodies by closing and restricting marine gateways. During the Miocene, for example, the Mediterranean basin became completely isolated from surrounding marine bodies in what is known as the ‘Messinian Salinity Crisis’. This period saw major changes to the regional water balance leading to evaporation and draw-down of the Mediterranean Sea. This had profound impacts on all aspects of the physical geography of the region including the climatology, biogeography, and geomorphology and its legacy can be seen across the region today. The more recent Quaternary geodynamics of the Mediterranean have generated an area which includes

a complex mixture of zones of plate subduction of various ages and stages (Figure 1.1b). The modern Mediterranean includes zones of active subduction associated with volcanic activity—such as the Calabrian arc—and older zones of now quiescent subduction such as the Betic-Rif arc. There is a wide range of seismic activity associated with these regions from deep (600 km) to shallow (1 km) and the development of extensive on-shore erosion (5.59–5.50 Ma) and deposition (5.50–5.33 Ma) of non-marine sediments in a lake sea (‘Lago Mare’). Krijgsman et al. (1999b) report that the event terminated with reflooding around 5.62–5.50 Ma and that the Mediterranean returned to normal marine conditions in the Pliocene around 5.33 Ma.

Causes It is generally agreed that glacio-eustatic fall (Adams et al. 1977; Hodell et al. 1986), horizontal crustal shortening (Weijermars 1988), tectonic uplift (Garcés et al. 1998; Hodell et al. 1989; Krijgsman et al. 1999a ) and climate change were the most likely triggers for the onset of the salinity crisis. Taking each of these mechanisms in turn we can attempt to ascertain their relative significance. With regard to a glacio-eustatic fall, the proposed 60 m drop in sea level in the Messinian is insufficient to have closed all the marine gateways. Also, the onset of evaporitic deposition at 05.96 ± 0.02 Ma

Tectonic Setting and Landscape Development

does not correspond with the benthic glacio-eustatic ‰18 O signature, suggesting that the Messinian Salinity Crisis is unlikely to be related to glacio-eustasy (Krijgsman et al. 1999b; Hodell et al. 2001). Palynological work in the Rifian corridor of northern Morocco (Warny et al. 2003) suggests that the regional climate was stable before, during, and after the Messinian Salinity Crisis. Crustal shortening and nappe emplacement is unlikely as it predates the Messinian Salinity Crisis (Comas et al. 1999). Thus the favoured mechanism is uplift. Duggen et al. (2003) account for the necessary uplift of c.1 km through the westward rollback of the subducted Tethys Oceanic lithosphere. With regard to the reflooding of the Mediterranean, data suggest that despite the Rifian corridor being the deepest of the oceanic gateways, reflooding probably occurred through the Gibraltar arc (Warny et al. 2003), creating a new marine gateway—the Strait of Gibraltar. The location of this gateway has been attributed to a mantle origin—involving gravity-induced slumping and faulting—by some workers (Duggen et al. 2003), perhaps assisted by geomorphic processes such as stream piracy (Blanc 2002). In the latter scenario an eastwardflowing stream that drained the eastern slope of an emergent Gibraltar Isthmus is thought to have breached the Atlantic/Mediterranean watershed. Blanc (2002) modelled the Atlantic inflow and suggests that the rate of Mediterranean recharge was exponential, with the level of the Mediterranean basin hardly changing in the first 26 years, but that it was completed within the ensuing 10–11 years. These dramatic and (on a geological timescale) extremely rapid changes in base level had profound consequences for the Messinian geomorphology of the region. The impacts are still visible in the regional Quaternary geomorphology.

Geomorphic Legacy The main geomorphic legacy of the Messinian Salinity Crisis relates to: (1) the dramatic changes in base level for the Mediterranean region; (2) the hydro- and erosional isostatic readjustments and (3) the widespread accumulation of evaporitic deposits and their associated distinctive karstic terrains. The dramatic lowering of base levels is credited with developing Mediterranean basin-wide subaerial erosion surfaces (Ryan 1978; Ryan and Cita 1978). The lowering of the Mediterranean Sea during the Messinian Salinity Crisis has been estimated at 500– 1,000 m (Mauffret 1979; Durand Delga 1980), 1500 m (Schlupp et al. 2001), 1,900–2,400 m (Gargani 2004), 2,500 m (Le Pichon et al. 1971; Ryan and Cita 1978;

15

Clauzon 1982) and 3,000 m (Malinerverno et al. 1981). However, the identification of the major erosion surface in some of the basins satellite to the main Mediterranean is hotly debated (see e.g. Fortuin et al. 2000 and Riding et al. 2000 and references therein). This erosion would have locally exacerbated any hydro-isotatic uplift as a result of erosional unloading (e.g. Gargani 2004). The presence of entrenched meanders and terraces within the Nile Canyon (formed when the Proto-Nile cut into the underlying Tortonian delta sediments as a result of the Messinian base-level fall) suggests sudden sporadic falls in sea level rather than a progressive drop (Barber 1981). Similarly, numerically modelled data from the Rhône suggests an initial drop in sea level of around 600–700 m followed by a secondary drop of 1,300–1,700 m. Gargani (2004) proposes that the first drop occurred over a period of 400 ka, and the second over a period of 50 ka. The associated loading and unloading of the Mediterranean basin water led to significant hydro-isostatic rebound. Two-dimensional flexure models (Norman and Chase 1983) suggest that the north-western and southeastern coasts of the Mediterranean would have generated shoreline bulges some 450 m high that were capable of reversing rivers with low stream power. Only those rivers with sufficient stream power (e.g. the Ebro, Po, Rhône, and Nile) could maintain their original courses through some of the areas of peripheral bulge. The hydro-isostasy would have been emphasized in the deepest parts of the Mediterranean basin such as the Ionian, Provençal, and central Tyrrhenian seas (Carminati and Doglioni 2004).

Mediterranean Rivers The main rivers of the Mediterranean, such as the lower Proto-Rhône and Ebro and Nile rivers, exhibit kmscale incision as a result of the dramatic lowering of base level in the Mediterranean during the Messinian (Clauzon et al. 1996). Outcrops and seismic data indicate deep palaeovalleys infilled with Pliocene sediments in the Rhône and Durance valleys (Audra et al. 2004). Canyon incision in the Aegean has been attributed to capturing the Black Sea drainage and leading to the deposition of ‘Largo Mare’ sediments (Hsü et al. 1973; McCulloch and De Deckker 1989). In response to the drop in base level a wave of incision up the main valley systems lead to headward erosion of valleys tributary to rivers such as the Po Canyon. These tributaries head-cut back into the Alpine highlands, forming the valleys later exploited by Quaternary glaciations. The use of boreholes and seismic mapping of the Nile Delta region (Barber 1981) shows that the basal

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Messinian erosion surface extends from approximately sea level in the southern delta area to depths of 5 km in the Nile cone. In the Nile region there is also subsurface evidence of the Messinian erosion surface that developed, and its associated drainage patterns. These appear to indicate that the lithology exerted a strong control on the development of the drainage pattern, with trellis drainage developed on resistant Oligocene basalts and dendritic drainage on less resistant Tortonian prodelta shales. The drop in base level of the Mediterranean led to river capture of the trellis drainage by the dendritic networks (Barber 1981) leading to sediment re-routing along parts of the Mediterranean shoreline.

Mediterranean Karst Karst landscapes within the Mediterranean region are commonly developed in Cretaceous limestones (Chapter 10), and more uniquely in the Messinian gypsum relating to the Messinian Salinity Crisis, and thus postdating it (e.g. Calaforra and Pulido-Bosch 2003; Ferrarese et al. 2003). The karsts developed within Cretaceous limestone often show a complex history that can be related back to the Mediterranean Salinity Crisis (Audra et al. 2004). Deep karst systems were probably formed during the Messinian Mediterranean drawdown, and reflooded during the Pliocene when the Mediterranean salinity crisis terminated. It is possible that palaeokarst systems such as the Rospo Mare oilfield (in Cretaceous limestone) located some 1,200 m below the modern Mediterranean sea level in the Adriatic (Soudet et al. 1994) may also owe their origin to this phase of the Mediterranean’s history. Off the coast of France the submarine spring of Port-Miou, south of Marseilles, is located in a drowned canyon of the Calanques Massif (Audra et al. 2004). Here the main water flow comes from a vertical shaft >147 m below present sea level. The shelf margin comprises a submarine karst plateau cut by a deep canyon, with the canyon reaching 1,000 m below sea level. In southern France inland karst systems are unusually deep, reaching >224 m below sea level at Fontaine de Vaucluse (Chapter 10). At the latter site flutes on the side of the vertical conduit, observed with a remote observation vehicle, indicate that this shaft was once air filled and connected to an underground river that was flowing towards a deep valley (Audra et al. 2004). Within the Ardeche a number of Vauclusian springs probably relate to the Messinian Rhône Canyon which is located 200 m below present sea level. Similarly the formation of karst systems in the Languedoc of southern France has also been linked with the drawdown of the Mediterranean during the salinity crisis (Josnin 2001). Josnin (2001) also suggests that

one of the primary controls on the karst in this region is the fracturing developed during the major tectonic phases of the Mediterranean. Within the latter systems reactivation of faults within the karst has led to the unblocking of abandoned conduits and their reincorporation into the active system. Mediterranean karst environments are discussed in detail in Chapter 10.

Geodynamics and the Nature of Quaternary Landforms It has already been shown that a combination of north– south compression and subduction rollback oblique/ orthogonal to this can explain the distribution of the main mountain ranges and basins of the Mediterranean region on geological (106 year) timescales. These mechanisms have provided the template for the landscape of the modern Mediterranean and can be used to explain the distribution and nature of key landforms.

Mountain Range Morphology The morphology of mountain ranges is largely controlled by tectonics, eustasy, climate, and lithology (e.g. Allen 1997; Leeder et al. 2002; Silva et al. 2003). Together these govern rates of erosion that, balanced against areas of crustal uplift, will lead to land elevation where erosion is less than uplift, i.e. surface uplift is occurring. Perhaps the most studied mountain range in the Mediterranean in this context are the Apennines of Italy. These are a fold and thrust belt that forms an accretionary wedge above the westward-directed Appenine– Maghrebides subduction zone (Lenci et al. 2004). Within mountain ranges the hydrological drainage divide and the highest elevations in mountain belts usually coincide, but they may move or grow as a function of the vergence and rate of the tectonic evolution (e.g. Ollier 1995). In the Appenines of Italy the most elevated topography is located eastward of the drainage divide. This has been attributed to lateral (eastward) migration of the topography associated with plate tectonic slab rollback (Figure 1.6). Lenci et al. (2004) used seismic lines to examine the relief and cross-sectional area of Italian mountain ranges. They found that in the Apennines where subduction depth was 200 km, and décollement depth (the depth of detachment of the upper cover from its substratum) was 10 km, the cross-sectional area of the mountains was 2000 km2 and relief greatest. Conversely, in the more southerly Calabrian arc, the subduction depth was 500 km, décollement depth was 3 km, the cross-sectional area of the mountains was

Tectonic Setting and Landscape Development

17

Denudation rate 2,000 km2 with a volume of 26 km3 . The total run out distance was some 110 km, with individual blocks up to 12,500 m × 3,000 m travelling some 3 km. It is considered that the likely triggering mechanism was a local earthquake event some 11,647–11,129 cal years BP (Canals et al. 2004).

Quaternary Landscape Development: Case Studies The main landform features discussed above will evolve in response to spatial and temporal changes in rates and styles of deformation. In addition the landforms do not act as separate entities but will interact to give an often complex response to ongoing tectonics. The developing landscape may reflect (1) direct tectonic control or (2) indirect tectonic control. Direct responses to tectonics occur in areas of rapid rates of deformation (e.g. the eastern Mediterranean). In the case of river systems this may lead to drainage diversion, for example reversed drainage. In areas of lesser tectonic activity (e.g. the western Mediterranean) the impacts of deformation may be less clear, but may be associated with equally dramatic river re-routing via river capture. Below, two case studies are used to illustrate the direct and indirect landscape responses to tectonics from two detailed case studies from the eastern and western Mediterranean. A recent special issue of Geomorphology edited by Silva et al. (2008) examines the impact of active tectonics on fluvial landscapes and drainage network development. With a strong focus on the Mediterranean, it includes case studies from the High Atlas of Morocco, the Dead Sea Rift, the Apennines of Italy, and various parts of the Iberian Peninsula.

Direct Impact of Tectonics: Southern Greece The direct impact of tectonics on the Quaternary landscape is perhaps best illustrated from the classic Gulf

21

of Corinth areas of southern Greece. Most of Greece is undergoing extension and subsidence as a result of subduction along the Hellenic arc, but areas such as northern Greece, where the Pindus Mountains are located, are undergoing long-term compression and uplift. Overall the landscape is made up of mainly east–west orientated grabens separated by faults from actively uplifting areas. This has created topographic features such as the Gulf of Corinth. Much of this active deformation relates to the motion between small plates around poles of rotation for each of these plates. In the central Mediterranean these plates are the Aegean, Ionian, and European plates and the poles are the Aegean–Ionian pole, the Aegean–European pole, and the Ionian–European pole (Figure 1.10). Of these boundaries the Aegean–Ionian boundary is typified by active subduction which varies from 3 cm a−1 in the north (Levkas) to 6 cm a−1 towards the south (Crete). The Aegean–European boundary is extensional with rates of movement of 2.5 cm a−1 in the west (Gulf of Corinth) to 5 cm a−1 in the east (Turkey). The relative poles of rotation generate these regional variations in rates of movement (Figure 1.10). The resultant tectonic style associated with these plates and their relative movements has uplifted mountainous areas such as Crete and Levkas (Chapter 16). The region is associated with large earthquakes on reverse

Fig. 1.10. Plate boundaries and the main poles of rotation referred to in the text. The poles presented here are those relevant to the motion at the plate boundaries of north-west Greece. The box indicates the location of the Gulf of Corinth and the image in Figure 1.11. Modified from King et al. (1997).

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faults in areas of compression and active seismicity and volcanism (Aegean Sea) in areas associated with the steeply (45◦ ) subducting Ionian Sea floor. The extension and compression which has affected the region has led to the development of the largest scale morphological features observable within the region, i.e. the basins and ranges. These features have often been inverted with former sediment receiving basins developed during extension becoming uplifted and exhumed source areas during compression. This section will now focus on one of these areas to examine the impact of these active tectonic processes on the associated geomorphological systems. The selected area is the welldocumented Gulf of Corinth and its eastern extent— namely the Alkyonides Gulf and inverted Megara basin (Figures 1.10 and 1.11).

Over the last 300 ka the uplift on the southern side of the Gulf of Corinth has been documented by the development of a number of erosional marine terraces formed during periods of high sea level (Kerauden and Sorel 1987; Armijo et al. 1996). The cumulative uplift for these terraces has been calculated at 1.5 mm a−1 but decreasing with time (Kerauden and Sorel 1987), and the relative displacement of blocks either side of a fault at 1 cm a−1 (King et al. 1997). These terraces are subject to active erosion and the rapid land degradation from gullying in this region can be clearly linked with active faulting (King et al. 1997). Over shorter timescales (individual earthquakes) significant deformation can occur rapidly. For example, in 1981 an earthquake of Ms 6.7 on 24 February affected the area of Corinth. Aftershocks of Ms 6.4 on 25 February and 4 March did further

Fig. 1.11. Oblique view of the Megara basin created with 3× vertical exaggeration using Geocover 2000 Landsat TM data courtesy of NASA World Wind. Note the drainage to the Alkyonides Gulf at the eastern end of the Gulf of Corinth cuts an erosional cirque into what is the footwall of the active Alepochori fault. Some 9 km3 of sediment has been removed by this drainage (current limit of the drainage divide is shown by the dashed line to the left of the image) over the last 1 Ma (Leeder et al. 1991). The drainage on the backtilted footwall shows tributaries dominantly draining towards the Megara fault, and the main axial drainages draining perpendicular to this trend, into the Saronic Gulf. The dotted line to the right of the image indicates the limit of stream incision on this surface. This asymmetric stream pattern reflects (1) an inherited tilt towards the inactive Megara fault, generating the early, north-east tributary drainage orientations, (2) backtilt from the active Saros fault, maintaining the NE/SW tributary drainage orientations and (3) backtilt from the active coastal scarp fault in the Gulf of Corinth generating the south-east orientated axial drainages.

Tectonic Setting and Landscape Development

damage. These led to significant geomorphological changes with some parts of the coast sinking by 1 m (Jackson et al. 1982; Vita-Finzi and King 1985; King et al. 1997; Chapter 16). Thus it is not uncommon to find juxtaposed areas of uplift and subsidence along some of the major active normal faults, with significant geomorphic implications (e.g. Leeder et al. 1991). The Alkyonides Gulf, at the eastern end of the Gulf of Corinth, is bounded to the south by a series of fault scarps, some of which were active during the 1981 earthquake (Jackson et al. 1982). Some of the individual fault scarps moved by as much as 1.5 m vertically, but averaged 0.5 to 0.6 m (Jackson et al. 1982) in the 1981 events. The topography associated with the main fault system descends to 400 m below sea level and the adjacent mountains are 1,000 m above sea level (Leeder et al. 1991). Indicators of active tectonics (hanging-wall subsidence) include drowned alluvial fans and associated marshes in the Psatha Bay beach zone (Leeder et al. 1991). These marsh areas are being progressively transgressed by backbeach washover sedimentation. The location of many marine (bays, lagoons, beach rock reefs, raised beaches) and continental (debris cones, alluvial fans) features is controlled by faulting and associated seismicity in this region (Leeder et al. 1991). Some of the debris cones, for example, are attributable to the 1981 sequence of earthquake events. Uplift in the footwall of the main Psatha-Skinos fault also controls stream incision, encouraging active headward erosion of stream networks (Leeder et al. 1991; Collier et al. 1993). Rates of drainage expansion and erosion within the area are mainly controlled by lithological variations associated with the faulting. The Megara basin of Greece is an inverted basin located at the eastern end of the Alkyonides Gulf. It is one of a number of such basins in this area. During the Neogene the basin was under extension and the bounding faults were tectonically active, with abundant evidence of syn-tectonic deformation affecting the contemporaneous sedimentary infill (more than 1 km of alluvial, fluvial, and lacustrine deposits; Bentham et al. 1991). In the Pleistocene (over about the last 1 Ma), the basin bounding faults became inactive and the basin was inverted due to uplift in the footwall of the active fault that bounds the south-east margin of the Alkyonides Gulf. The basin fill was then subjected to weathering, erosion, and the development of drainage networks. Some 9 km3 of footwall sediments have been eroded from the backtilted footwall of the main fault scarp (Figure 1.11). This gives a mean erosion rate of 0.27 mm a−1 and the volume would be sufficient to deposit a layer 180 m thick over the Alkyonides

23

Gulf (some 50 km2 ; Leeder et al. 1991). These rates of erosion are strongly aided by the availability of highly erodible (Plio-Pleistocene basin fill) lithologies. The change in fault activity was probably related to clockwise crustal block rotations of some 25–3◦ that occurred about this time, causing reorientation of the faults from east to west to 115–12◦ (Leeder et al. 1991). This change in tectonic activity was associated with the initiation of new drainage systems and a change in basin slope. Previously the Megara basin general palaeoslope was towards the north-east. Since then it has been back-tilted towards the south-east. This change in slope is recorded in the drainage pattern. The highest order, axial drainages drain to the south-east down the current dip slope (Figure 1.11). The tributaries to these streams are dominantly located to the south-west and flow towards the north-east, suggesting they may be streams originally developed on the northeast dipping palaeoslope (Goldsworthy and Jackson 2000).

Indirect Impact of Tectonics: South East Spain Although many long-term landscape elements of the Mediterranean can be attributed directly to the geodynamics of the region, others are less direct and indicate the complex response of geomorphic systems to external controls. Much of this will be examined in detail in later chapters of this book. Here we focus on a specific process which is abundant throughout the Mediterranean as a function of the regional tectonics—river capture. River capture in its true sense (a bottom-up process) is prolific in areas with high relief, a mixture of dip and strike drainages, and differential uplift. These characteristics provide optimal conditions for river capture to occur i.e. where a lower elevation stream undergoing headward erosion and gullying can breach the drainage divide of a higher elevation drainage system and lead to stream capture. River capture has occurred in response to the Mediterranean drawdown in the Messinian (e.g. Barber 1981) and has even been credited with facilitating the reflooding of the Mediterranean from the Black Sea (Hsü et al. 1973; McCulloch and De Deckker 1989) and Atlantic (Blanc 2002). River capture can lead to sudden, dramatic falls in base level within river systems (for example some 500 m in the Guadix basin, south-east Spain; Calvache and Viseras 1997) with major impacts on sediment flux and routing both within and between drainage systems (e.g. Mather and Harvey 1995; Mather 2000a , b; Mather et al. 2000, 2002; Stokes et al. 2002; Azañón et al. 2005).

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Anne Mather

River Watershed, Río de Aguas Watershed, Rambla de los Feos R a G mb óc la ha de r

R Cin ambl ta a d Bla e nc a

Rambla de Sorbas/ Río de Aguas

los de bla Ram hopos C

Ra

Badlands

de la mb eos F

Landslides

los

>105 m3 >103 – 105 m3 tenfold increase in incision was experienced (Stokes et al. 2002) and this radically altered the sediment delivery processes (Figure 1.14). Initially the generation of steep, rapidly unloaded slopes generated mass movement processes such as landslide

failures. In weaker lithologies the dominant sediment delivery process later became dominated by more progressive slope erosion by surface runoff and subsurface piping processes (pseudokarst). Most of this accelerated erosion is still restricted to the main valley sideslopes, and has not yet reached the main drainage divides so that overall surface lowering is much less than that recorded by the localized valley incision (Figure 1.15). Post-capture channel incision was 50 m about 7 km upstream of the capture site and 25 m about 13 km upstream. In these areas a complex reorganization of tributary drainages through stream capture occurred (Mather 2000a) resulting in extensive badland erosion and the development of landslides of 290,000 m3 to 550,000 m3 in volume (Mather et al. 2003).

Conclusions To understand the individual landscape elements of the Mediterranean basin it is evident that we have to understand the geodynamics of the region on a variety of spatial and temporal scales. Whether examining the long-term evolution of a badland, a karst system, or a river catchment we need to appreciate that many aspects may be inherited from earlier periods of tectonic activity and landscape change such as the Messinian Salinity Crisis. Individual seismic events can lead to catchmentwide landsliding, generating sediment sources for river systems that may take many millennia to be eroded 100

13

80 60

R

12

bl

a

de

40

G

óc

11

Surface lowering (m)

20

ha

r 0 as/ orb e S uas g la d Rb per A Up

10 09 08

R Cin bla d ta Bla e nc a

Modern drainage network

07 06

S

e la d s Rb hopo C los

05 73

74

75

76

77

78

79

80

81

Fig. 1.15. Surface lowering above the 70 ka river capture site depicted in Figure 1.12. Note the highest levels of surface lowering are associated with (1) the valley networks, (2) valley confluences, and overall with (3) the lower reaches of the drainage network proximal to the capture point (which is just off the map to the right). Modified from Mather et al. (2002).

Tectonic Setting and Landscape Development

and transported through the fluvial networks. We also need to appreciate that whilst many processes may be directly attributable to the regional tectonics, such as drainage reversal, the indirect impact of tectonics is not always immediately clear. Indirect impacts, such as river capture, can lead to sudden, high-magnitude changes in geomorphic systems that are still felt in the landscape today, despite (in human terms) their relative antiquity.

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Blanc, P.-L. (2002), The opening of the Plio-Quaternary Gibraltar Strait: assessing the size of a cataclysm. Geodinamica Acta 15: 303–17. Briole, P., Rigo, A., Lyon-Caen, H., Ruegg, J. C., Papazissi, K., Mitsakaki, C., Balodimou, A., Veis, G., Hatzfield, D., and Deschamps, A. (2000), Active deformation of the Corinth rift, Greece: Results from repeated Global Positioning System surveys between 1990 and 1995. Journal of Geophysical Research 105: 605–25. Buforn, E., Udias, A., and Madariaga, R. (1991), Intermediate and deep earthquakes in Spain. Pure and Applied Geophysics 136: 375–93. Calaforra, J. M. and Pulido-Bosch, A. (2003), Evolution of the gypsum karst of Sorbas (SE Spain). Geomorphology 50 (1–3): 173–80. Calais, E., DeMets, C., and Nocquet, J. M. (2003), Evidence for a post- 3.16 Ma change in the Nubia-Eurasia plate motion. Earth and Planetary Science Letters 216: 81–92. Calcaterra, D. and Santo, A. (2004), The January 10, 1997 Pozzano landslide, Sorrento Peninsula, Italy. Engineering Geology 75: 181–200. Calvache, M. L. and Viseras, C. (1997), Long-term control mechanisms of stream piracy processes in southeast Spain. Earth Surface Processes and Landforms 22: 93–105. Campbell, I. A. (1989), Badlands and badland gullies, in D.S.G. Thomas (ed.) Arid Zone Geomorphology. Belhaven, London, 158–83. Canals, M., Lastras, G., Urgeles, R., Casamor, J. L., Mienert, J., Cattaneo, A., De Batist, M., Haflidason, H., Imbo, Y., Laberg, J. S., Locat, J., Long, D., Longva, O., Masson, D. G., Sultan, N., Trincardi, F., and Bryn, P. (2004), Slope failure dynamics and impacts from seafloor and shallow sub-seafloor geophysical data: case studies from the COSTA project. Marine Geology 213: 9–72. Candy, I., Black, S., and Sellwood, B. W. (2005), U-Series isochron dating of immature and mature calcretes as a basis for constructing Quaternary landform chronologies for the Sorbas Basin, southeast Spain. Quaternary Research 64: 100–11. Carey, S. W. (1955), The orocline concept in geotectonics. Papers and Proceedings of the Royal Society of Tasmania 89: 255–88. Carminati, E. and Doglioni, C. (2004), Mediterranean Tectonics. Encyclopaedia of Geology, Elsevier, Amsterdam. Carro, M., De Amicis, M., Luzi, L., and Marzrati, S. (2003), The application of predictive modelling techniques to landslides induced by earthquakes: the case study of the 26 September 1997 Umbria-Marche earthquake (Italy). Engineering Geology 69: 139–59. Chester, D. K. and Duncan, A. M. (1982), The interaction of volcanic activity in Quaternary times upon the evolution of the Alcantara and Simento Rivers, Mount Etna, Sicily. Catena 9: 319–42. Clauzon, G. (1982) The Messinian Rhône canyon as a definite proof of the “desiccated deep-basin model”. Bulletin Société Géologique de France 24: 597–610. Suc, J.-P., Gautier, F., Berger, A., and Loutre, M.-F. (1996), Alternate interpretation of the Messinian Salinity Crisis: Controversy resolved? Geology 24: 363–6. Collier, R. E. L., Leeder, M. R., and Jackson, J. A. (1995), Quaternary drainage development, sediment fluxes and extensional tectonics in Greece, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Balkema, Rotterdam, 31–44.

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2

The Marine Environment: Present and Past Eelco Rohling, Ramadan Abu-Zied, James Casford, Angela Hayes, and Babette Hoogakker

Introduction The Mediterranean is a landlocked, semi-enclosed marginal sea that spans a maximum of 3,860 km in the west–east direction, and a maximum of ∼1, 600 km in the north–south direction. Along its roughly 46,000 km of coastline, the basin is enclosed by mountainous terrain, except for a part of the North African margin to the east of Tunisia. The Mediterranean Sea contains very deep basins, more than 4 km, and has an average depth of approximately 1,500 m. Its only natural connection with the open (Atlantic) ocean is through the narrow Strait of Gibraltar, which contains a 284-m deep sill (at a width of ∼30 km), and reaches a minimum width of only 14 km (at a depth of 880 m) (Bryden and Kinder 1991). The Strait of Sicily subdivides the Mediterranean Sea into a western and an eastern basin. This strait is relatively wide (about 130 km) and contains a topographically complex sill-structure with an estimated average depth of 330 m (Wüst 1961), reaching 365 and 430 m in the two major channels (Garzoli and Maillard 1979). The eastern Mediterranean contains two smaller marginal basins, namely the Adriatic Sea and the Aegean Sea (Figure 2.1). Watermasses are exchanged through both the Strait of Gibraltar and the Strait of Sicily by eastward surface and westward subsurface flows (Figure 2.2). This pattern of exchange results from a net buoyancy loss in the basins on the easterly side of the sills, primarily due to strong net evaporative loss from the Mediterranean, and secondarily to some net cooling. Deep

water ventilation in the Mediterranean is primarily salt-driven, and secondarily temperature-driven. This is similar to the mode observed in the present-day Red Sea, but contrasts with the temperature-dominated mode in the modern world ocean. As such, the Mediterranean deep ventilation might be more appropriately described as halo-thermal rather than with the common term thermo-haline. This offers a useful analogue for world ocean circulation modes in past times with very warm and relatively equable global climates, such as the Mesozoic. Interestingly, the Mediterranean is characterized by periodic, widespread deposition of organic-rich sediments or ‘sapropels’ over periods of several thousands of years, similar (in miniature) to the deposition of ‘black shales’ in the Mesozoic oceans. Surface water flowing in through the Strait of Gibraltar is traceable through the Strait of Sicily into the eastern Mediterranean, although its salinity increases steadily towards the east (e.g. Wüst 1961; MalanotteRizzoli and Hecht 1988; Malanotte-Rizzoli and Bergamasco 1989; Pinardi and Masetti 2000) (Figures 2.2 and 2.3). The eastward salinity increase culminates in values around 39.2 psu (up to an extreme of 39.5 psu, Wüst 1960) in the eastern Levantine sector of the Mediterranean, compared with 36.1–36.2 psu for the Atlantic inflow at Gibraltar. The high Levantine salinities are associated with high temperatures in summer, but strong winter cooling (especially between Cyprus and Rhodes) causes surface waters to attain high enough densities to sink and spread at intermediate depths (150–600 m). This forms the

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Eelco Rohling et al.

Fig. 2.1. Map of the Mediterranean Sea.

Fig. 2.2. Longitudinal cross-section showing water mass circulation in the Mediterranean Sea during the present-day winter (modified from Wüst 1961). Isolines indicate salinity values in psu (practical salinity units) and arrows indicate the direction of water circulation in the Mediterranean Sea.

‘Levantine Intermediate Water (LIW)’. This watermass spreads westward from its formation area throughout the entire Mediterranean Sea. Admixtures of regional winter mixed-layer waters slightly reduce the LIW salinity as it spreads, transforming this watermass into what has become known as ‘Mediterranean Intermediate

Water (MIW)’. There are also contributions of Eastern Mediterranean Deep Water (EMDW) and Western Mediterranean Deep Water (WMDW) to the MIW upon its passage through the Strait of Sicily and the Strait of Gibraltar. In most parts of the eastern Mediterranean, MIW salinities are between 38.8 and 39.1 psu, while

The Marine Environment: Present and Past

35

Fig. 2.3. Surface water circulation in the Mediterranean Sea (modified from Vergnaud-Grazzini et al. 1988; Roussenov et al. 1995). Shaded areas indicate intermediate and deep water formation.

values in the western Mediterranean are between 38.5 and 38.8 psu. The subsurface outflow from the Mediterranean through the Strait of Gibraltar has a salinity of 38.2–38.4 psu (among many others: Wüst 1960, 1961; Garzoli and Maillard 1979; Gascard and Richez 1985; Bryden et al. 1994). The influence of Mediterranean Outflow can be traced as a salinity maximum centred on about 1,000 m depth in the North Atlantic (e.g. Reid 1979; Hill and Mitchelson-Jacob 1993; Iorga and Lozier 1999, O’NeillBaringer and Price 1999). This maximum represents the overall average signature, but an important component of the dispersal of Mediterranean Outflow within the North Atlantic has been found to occur in the form of discrete subsurface ‘lenses’ of salty and warm Mediterranean water. These are the so-called Mediterranean eddies or ‘Meddies’ with diameters up to 100 km, the pathways of which have been traced with neutralbuoyancy floats (Richardson et al. 1991, 2000). The isopycnals (lines of equal water density) at which Mediterranean outflow settles show northward shoaling within the north-east Atlantic. Near the Iceland– Scotland Ridge deep winter mixing of fresher surface waters with the salty Mediterranean tongue raises the salinity of the surface water that enters the Norwegian Sea through the Faroe–Shetland Channel (Hill and

Mitchelson-Jacob 1993). This preconditions the inflow for later convection by increasing its salinity by several tenths of a psu relative to ‘background’ (Reid 1979), the density equivalent of 1–2◦ C cooling. Such preconditioning may facilitate the formation of North Atlantic Deep Water (NADW) in the Norwegian Sea (Reid 1979; Hill and Mitchelson-Jacob 1993). Returning attention to the Mediterranean now, EMDW and WMDW are found below about 1 km depth in the eastern and western Mediterranean basins, respectively, separated by the sill in the Strait of Sicily. Between about 600 and about 1,000 m, a transitional watermass is found between the deep waters and MIW. WMDW is formed in the northern sector of the western Mediterranean, notably in the Gulf of Lions, due to strong winter cooling caused by cold continental air outbreaks that are orographically channelled towards the basin via the Rhône valley (the ‘Mistral’). EMDW is formed in two separate regions, namely the Adriatic Sea and the Aegean Sea. Both areas are subject to orographically channelled continental air outbursts in winter, the ‘Bora’ over the Adriatic, and the ‘Vardar’ over the Aegean Sea (Chapter 3). In schematic terms, the Mediterranean deep water ventilation can be viewed as a two-stage motor (a detailed explanation is given below). The first stage

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consists of the dominantly salt-driven formation of LIW. The salt distributed throughout the Mediterranean Sea by LIW/MIW preconditions the basin for deep water formation. The second stage of the deep-ventilation motor is dominated by cooling events related to orographically channelled continental air outbursts over northern sectors of the basin in winter. Given the presence of a major monsoon-fed river (Nile) in close proximity to the centre of action of the salt-driven first stage of the motor, we can expect that monsoon variations would be reflected in the efficiency of the Mediterranean’s deep ventilation, as deep water preconditioning would be directly affected. However, we can also expect important northerly climate impacts on the deep ventilation, related to changes in the frequency and/or intensity of winter cooling events. The small volume of the Mediterranean Sea, compared with ocean basins, causes changes in its climatic forcing to be recorded virtually instantaneously in palaeoceanographic proxy data, such as stable isotope and other geochemical ratios, and microfossil abundances. The basin’s limited communication with the open ocean implies that any climatic signals will be recorded in an amplified fashion in Mediterranean properties, such as salinity and specific elemental concentrations. The critical location of the Mediterranean Sea on the boundary between a subtropical/monsoon regime and the temperate westerlies means that it is highly sensitive to changes in both these systems. Since both systems primarily affect fundamentally different characteristics of the basin, the Mediterranean is an excellent site for study of the relative timing and impact of changes in the two major systems (subtropical/monsoon climate predominantly affects freshwater balance, while the temperate westerly climate controls cooling in the north). The regularly recurring deposition of organic-rich ‘anoxic’ sapropels offers discrete windows for very high-resolution study, since these intervals are not affected by sediment homogenization due to bioturbation. The Mediterranean Sea may hold important clues as to the functioning of circulation in the Mesozoic oceans and the formation processes of black shales that are of great economic importance. To appreciate fully the processes underlying past changes in Mediterranean climate and hydrography, comprehensive background knowledge of present-day conditions is indispensable. Following a brief history of the development of the Mediterranean basin, therefore, this chapter first discusses relevant aspects of the region’s modern climate and oceanography before

engaging in a review of palaeoclimatological and palaeoceanographic reconstructions.

Long-term Context Chapter 1 of this volume deals with the long-term tectonic history of the Mediterranean basin, and this section therefore only highlights several particularly relevant aspects of change in the climatic and oceanographic setting. Overall, the Mediterranean is a relic ocean basin, representing the final stage of closure of the Tethys Ocean prior to continent–continent collision as the African plate converges with the Eurasian plate system. Note, however, that parts of the western basin are relatively young, and actively opening and deepening— notably the Tyrrhenian Sea. The proto-Mediterranean’s eastern connection with the open ocean (through the Levantine–Arabian region) closed roughly 18 Ma ago (Vergnaud-Grazzini 1985). Since that time, the only connection of the Mediterranean basin with the open world ocean has been through waterways in the west. Two such waterways connected the Mediterranean with the Atlantic Ocean: one across north Morocco (Rifian Strait) and one through the southern Iberian Peninsula (Betic Strait). Tectonic closure of the Betic and Rifian Straits led to massive evaporite deposition between 5.9 and 5.5 Ma, a phase known as the ‘Messinian Salinity Crisis’ (followed by the so-called ‘Lago Mare’ phase 5.5– 5.3 Ma) (Hilgen et al. 1995; Chapter 1). Re-establishment of open marine conditions following the Messinian Salinity Crisis appears to have been virtually synchronous everywhere in the Mediterranean basin, and may be ascribed to the tectonic opening of the Strait of Gibraltar. This event heralded the appearance of a basin that clearly began to approach the modern configuration. However, ongoing plate subduction processes (including slab detachment underneath southern Italy) caused continuing highly complex tectonic reshaping in the area. For example, the Tyrrhenian Sea underwent very rapid deepening and extension between ∼3 and ∼1.5 Ma, while tremendous uplift in southern Italy and parts of Greece has caused Late Pliocene/Early Pleistocene coastal sediments to be displaced to many hundreds of metres above modern sea level. The basin’s geological history can directly affect modern processes. When faults expose parts of the massive Messinian evaporite deposits to sea water in the basin, salt dissolution affects modern bottom-water properties.

The Marine Environment: Present and Past

This happens in the so-called ‘brine basins’ of the eastern Mediterranean. Dissolved salts in the bottom waters of these isolated depressions cause extremely high salinities, separated from the normal deep waters by a very sharp salinity gradient (halocline), which defines a strong density gradient (pycnocline). The oceanography and chemistry of brine basins are entirely different than in the open waters around them. Because of the extreme density stratification, the brines are not ventilated, and thus have become entirely oxygen-depleted. Sediments in these basins are often disturbed by masstransport processes, but on rare occasions undisturbed sections yield beautifully laminated cores, reflecting the fact that there is no benthic life to homogenize the sediments through bioturbation (among many others: Jongsma et al. 1983; Scientific Staff Cruise BAN84 1985; Troelstra et al. 1987; MEDRIFF consortium 1995; Wallmann et al. 1997). Major ‘global’ climate developments also need to be considered when studying palaeoclimatic and palaeoceanographic signals in the Mediterranean. The development towards a ‘glacial mode’ in the Northern Hemisphere started around 3.2–3.1 Ma (Shackleton and Opdyke 1977; Thunell and Williams 1983; Prell 1984). The Mediterranean environment was substantially affected by the Northern Hemisphere glaciations (Vergnaud-Grazzini 1985; Thunell et al. 1987, 1991). Ruddiman et al. (1987) found the first clear evidence for ice-rafting in the North Atlantic around 2.55 Ma, and Zachariasse and Spaak (1983) demonstrated that biogeographic patterns similar to the present originated around that time in the Mediterranean and adjacent Atlantic. The early development of Northern Hemisphere glaciation was associated with climatic change over the Mediterranean basin, characterized by increasing seasonal contrasts with very dry summers (Suc 1984; see also Thunell 1986). Suc (1984) argued that the ‘modern’ conditions with cool wet winters and hot dry summers first developed around 3.2 Ma, and that summer drought became more persistent after 2.8 Ma. Global atmospheric circulation modelling by Ruddiman and Kutzbach (1989) suggests that these developments may have resulted from northern hemispheric climate reorganization due to uplift of the Tibetan plateau, while the periodical appearance of steppe vegetation in the Mediterranean realm since 2.3 Ma (Suc 1984) would be related to large-scale expansions of Northern Hemisphere ice-sheets. The early glacial cycles had a mean periodicity of 41,000 years (obliquity forcing), which changed to a predominant periodicity of 100,000 years (eccentricity forcing) after the so-called

37

Mid-Pleistocene Transition, roughly 1.0 to 0.9 Ma (Shackleton and Opdyke 1973, 1976; Pisias and Moore 1981; Ruddiman et al. 1986, 1989). This change is well represented in Mediterranean isotopic, floral, and faunal records (e.g. Zachariasse et al. 1989, 1990; Lourens et al. 1992; Vergnaud-Grazzini et al. 1993; Chapter 4).

Modern Climate and Oceanography Climate The classical Mediterranean climate, discussed in detail in the following chapter, is characterized by warm and dry summers, and mild and wet winters. As such, it appears opposite to monsoon climates, which instead comprise a pluvial maximum in the warm months. The Mediterranean climate regime is due to the basin’s location on the transition between the climate conditions of the temperate westerlies that dominate over central and northern parts of Europe, and the subtropical high pressure belt over North Africa (Figure 2.4) (Boucher 1975; Lolis et al. 2002; Chapter 3). In summer, the subtropical high pressure conditions are displaced to the north and most of the Mediterranean experiences drought, especially the south-eastern sector. Polar front depressions may still reach the western Mediterranean, but they only exceptionally penetrate the eastern Mediterranean (Rohling and Hilgen 1991). During winter, the subtropical conditions are displaced southward, and the (northern sector of) the Mediterranean comes under the influence of the temperate westerlies with the associated Atlantic depressions that track eastward over Europe. Polar and continental air masses over Europe are channelled towards the Mediterranean through valleys between the mountainous topography of the northern Mediterranean margin. During winter and spring, intense cold and dry katabatic air flows are channelled through the lower Rhône Valley towards the Gulf of Lions (‘Mistral’), and also over the Adriatic and Aegean Seas (‘Bora’ and ‘Vardar’), causing strong evaporation and cooling of the sea surface (e.g. Leaman and Schott 1991; Saaroni et al. 1996; Poulos et al. 1997; Maheras et al. 1999; Casford et al. 2003; and references therein). Conditions for northerly air flow into the western and eastern Mediterranean are determined by interaction between an intense low over the central or eastern Mediterranean, and north-eastward extension of the Azores High (over Iberia, France, and

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Eelco Rohling et al.

North Hadley Cell 23°N

L

ITCZ 0°

Tropical easterly jet South Hadley Cell

H 23°S

Ferrell Cell

Fig. 2.4. Northern Hemisphere summer atmospheric circulation pattern. The main winds are indicated as arrows. ITCZ = Inter-Tropical Convergence Zone, H = areas of high sea-level pressure, L = areas of low sea-level pressure (modified from Rossignol-Strick 1985; Reichart 1997).

southern Britain) or westward ridging of the Siberian High towards north-west Europe and southern Scandinavia (Maheras et al. 1999; Lolis et al. 2002). The wintertime low surface pressure conditions over the Mediterranean are a direct consequence of the high sea-surface temperatures due to the high thermal capacity of the basin’s watermasses (Lolis et al. 2002). The most pronounced basin-wide cold winter events (complementing widespread cold conditions over Europe) develop in association with positive sea-level pressure anomalies to the west or north-west of the British Isles and particularly low pressure over the Mediterranean, a configuration that reflects an extreme phase of the North Atlantic Oscillation (NAO) (Moses et al. 1987; Maheras et al. 1999). The main mode of climate variability in the Mediterranean is expressed by the so-called Mediterranean Oscillation (MO), a west–east pressure see-saw that is apparent both at the surface and at 500 hPa, especially in winter and spring (Maheras et al. 1999; Lolis et al. 2002). Statistical correlation has been found between the MO and the pressure see-saw of the NAO, where the low NAO index phase is associated with wet conditions in the western Mediterranean (Maheras et al. 1999; Lolis et al. 2002; Dünkeloh and Jacobeit 2003; and references therein). This confirms previous observations of direct NAO impacts on the western

Mediterranean, but for the eastern basin the relationship remains weakly established, except via dependence of the MO on the NAO (Dünkeloh and Jacobeit 2003). The statistically second main mode of winter variability, with important impacts on cyclogenesis in the basin and consequent precipitation in the north-eastern and southcentral parts of the Mediterranean, is the Mediterranean Meridional Circulation (MMC) (Dünkeloh and Jacobeit 2003). Cold and relatively dry northerly (meridional) airflow over warm sea surfaces causes intense cyclogenesis (formation of new depressions) in the northern sectors of the Mediterranean. Most cyclones observed in the Mediterranean are thus formed over the basin itself, although some Atlantic depressions may enter the (western) basin (Rumney 1968; Trigo et al. 1999; Chapter 3). Throughout the basin, however, winter cyclones are clearly linked to North Atlantic synoptic systems, as secondary lows when Atlantic systems interact with the Alps with additional cyclogenesis over the basin itself (Trigo et al. 2000). Cyclogenesis is most frequent in the western Mediterranean, especially over the Gulf of Genoa and Ligurian Sea, but the Aegean Sea is a major centre for winter-time cyclogenesis as well (Trewartha 1966; Rumney 1968; Boucher 1975; Cantu 1977; Trigo et al. 1999; Chapter 18). Most of the Genoan depressions track south-eastward down the coast of Italy

The Marine Environment: Present and Past

and then eastward or north-eastward across the Aegean Sea or northern Levantine seas (Trewartha 1966; Rumney 1968; Trigo et al. 1999; Lolis et al. 2002). Along the way, these depressions as well as those developing over other centres of cyclogenesis cause the winter precipitation that is so typical for the modern Mediterranean climate. The stable hydrogen and oxygen isotope composition of this precipitation follows a Mediterraneanspecific mixing line (the Mediterranean Meteoric Water Line, MMWL), which is different from the global Meteoric Water Line (MWL) due to the dominant contribution of moisture evaporated from the Mediterranean Sea into low-humidity air masses (Matthews et al. 2000; and references therein). Summer rainfall is low today, especially in the eastern basin. Although cyclogenesis occurs around Cyprus and the Middle East in summer, as a semi-permanent extension of the Indian monsoon low, dry summer conditions prevail as a consequence of adiabatic descent in the upper troposphere that is related to the intense Asian summer monsoon (Rodwell and Hoskins 1996; Trigo et al. 1999). Mean annual precipitation along the Mediterranean ranges from less than 120 mm in North Africa, to over 2,000 mm in portions of south-west Turkey and in the eastern Adriatic Sea along the slopes of the Dinaric Alps (Naval Oceanography Command 1987). Total evaporation in the entire Mediterranean increases towards the east, with an average of 1,450 mm y−1 (Malanotte-Rizzoli and Bergamasco 1991) to 1,570 mm y−1 (Béthoux and Gentili 1994). Strong rates of evaporation occur in areas subjected to strong winds, such as the Gulf of Lions and Ligurian Sea, the Aegean and Cretan Seas, and the southern part of the Turkish coast (MEDOC Group 1970; Miller 1974). Evaporation is weakest along the Moroccan and Algerian coasts (The Alboran Sea) where the air masses generally arrive from the Atlantic with relatively high air humidity. The basin-wide mean Mediterranean excess of evaporation over freshwater input [E (evaporation) – P (precipitation) – R (runoff)] has been variously estimated at ∼1,000 mm y−1 (Béthoux et al. 1999), 750 mm y−1 (Gilman and Garrett 1994), and 560–660 mm y−1 (Bryden and Kinder 1991). There is marked spatial variation in regional values (Chapters 3 and 8). Northern areas such as the Gulf of Lions, Adriatic and Aegean Seas show relatively low excess evaporation rates due to high freshwater inputs from the Rhône and Ebro rivers, the Po River, and the Black Sea, respectively. Southern regions show very high excess evaporation rates, especially in the eastern Mediterranean (Béthoux and Gentili 1994). The strong overall excess

39

evaporation results in a pronounced surface water salinity increase from west to east (MEDATLAS 1997) (Figure 2.2). Sea surface temperature values in the Mediterranean reflect a balance dominated by high energy gain from solar irradiation during the widespread subtropical high-pressure (clear) conditions in summer, and considerable (latent) heat loss during evaporation. As a result, sea surface temperature values increase towards the east and south throughout the Mediterranean. Winter values are around 10◦ C in the north-western Mediterranean and 15◦ C in the south-eastern Mediterranean, while summer values are around 21◦ C in the northwestern Mediterranean and 26◦ C in the south-eastern Mediterranean (Naval Oceanography Command 1987). The warmest season centres on July–August and the coldest on February–March. One further climate impact on the Mediterranean Sea must be mentioned. It concerns a ‘remote’ influence by a climate system that does not itself penetrate into the basin, namely the African monsoon. It used to influence the Mediterranean mainly through Nile River discharge, but has been severely curtailed since completion of the first stage of the Aswan High Dam in 1964. Prior to the anthropogenic control of the Nile, its average discharge was 8.4 × 1010 m3 yr−1 (4.5 × 1010 m3 yr−1 in lowflood years to 15.0 × 1010 m3 yr−1 in high-flood years), which from the mid-1960s has dwindled to a negligible amount (Nof 1979; Said 1981; Wahbi and Bishara 1981; Béthoux 1984; Rohling and Bryden 1992). Note that the reported discharge values illustrate that, even in the instrumental era, there was strong (threefold) interannual variability between high and low discharge years, which was mainly related to variability in the monsoon-fed contribution of the Blue Nile and Atbara rivers (see below). The Nile River comprises two different systems: the White Nile, which drains the equatorial uplands of Uganda in a regular, permanent manner; and the Blue Nile and Atbara, which receive highly seasonal African monsoon precipitation from the Ethiopian highlands. Nile hydrology has been summarized by Adamson et al. (1980) and Williams et al. (2000). In summary, these authors find that prior to extensive anthropogenic intervention (damming), a maximum of 30 per cent of the annual discharge of the Nile originated from the White Nile, and a minimum of 70 per cent from the Blue Nile and Atbara basins. The winter flow was dominated (83%) by the steady White Nile contribution, whereas the Blue Nile/Atbara component provides 90 per cent of the flow in summer. This seasonal contrast results from a massive increase in the Blue Nile and

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TABLE 2.1. Contributions to total Nile discharge from the main tributaries Flood season (between August and October). Values in 106 m3 day−1

Regular flowing of water (outside flood seasons). Values in 106 m3 day−1

White Nile (from Lake Victoria = equatorial highlands) Blue Nile (Ethiopia) Atbara (Ethiopia)

70.0

37.5

485.0 157.0

7.5 0.0

Total

712.0

45.0

Sources: Hurst (1944); Said (1981).

Atbara discharge between a winter low and summer high (see also Table 2.1), with the monsoon-related peak occurring in the months August–October. The White Nile discharge shows a much smaller ratio of change between its annual peak and lowest monthly value (Table 2.1), and is highest between late September and January. Table 2.1 illustrates historical discharge values for the three main tributaries after Hurst (1944) and Said (1981) (according to those observations, the White Nile contribution to total annual discharge amounts to only 14%). The total suspended sediment load transport to the Mediterranean coast before closure of the Nile by the Aswan High Dam exceeded 1.0 × 108 tonnes yr−1 (Sharaf el Din 1977; El Dardir 1994; Stanley 1996). Since completion of the Aswan High Dam, there has been negligible Nile discharge and sediment transport into the Mediterranean through the Rosetta and Damietta outlets (UNDP/UNESCO 1978). Instead, salt water entering the mouth of the Rosetta extends some 25 km upstream to the Nile barrage at Mutubis. A little fresh water reaches the Mediterranean through the Manzalla, Burullus, and Idku lagoon outlets, and by pumping of lake Maryut water to the sea at Alexanderia (Stanley and Wingerath 1996). The suspended load that bypasses the Nile Delta to the shelf via Nile distributaries, lagoon outlets, and canals is about 15 per cent of the original (predam) load (Stanley et al. 1998). Apart from damming, the freshwater flow and sediment flux into the Mediterranean Sea were also curtailed due to the extensive irrigation network of canals and drains covering the entire Nile delta. Prior to its anthropogenic reduction, the Nile plume used to be distinctly traceable with the prevailing surface circulation in the easternmost Mediterranean, from the Nile delta east- and northwards along the eastern Levantine coast. It caused a zone with notably reduced

surface-water salinities and enhanced turbidity (suspended matter) (Reiss et al. 1999).

Surface-water Circulation Circulation in the Mediterranean Sea is driven by wind stress and thermohaline forcing (Robinson et al. 1992). Atlantic water (AW) enters the Mediterranean Sea as a surface flow through the Strait of Gibraltar, compensating for the net evaporative loss from the basin and the subsurface outflow. AW enters with a salinity of about 36.2 psu and temperature of about 15◦ C (Béthoux and Gentili 1994). As it migrates through the Strait of Gibraltar, AW mixes with upwelled Mediterranean Intermediate Water (MIW), creating Modified Atlantic Water (MAW) which has higher temperatures (16◦ C) and salinities (36.5%) (La Violette 1986; Tintoré et al. 1988; Arnone et al. 1990; Heburn and La Violette 1990). In the Alboran Sea, MAW is present along the southern Spanish coast as a strong jet (with speeds up to several kilometres an hour) approximately 20 km wide and extending to a depth of 150 m (Pistek et al. 1985). The strength of the jet initiates the formation of two anticyclonic gyres (Figure 2.5), the positions of which fluctuate on timescales of 3–4 weeks (Heburn and La Violette 1990). As the MAW flows eastward along the Spanish coast to Almeria, it converges with resident Mediterranean waters. The subsequent deflection of MAW towards Oran on the Algerian coast forms a well-defined frontal zone along the eastern edge of the Eastern Alboran Gyre (Figure 2.5). This front extends to a depth of 200 m and has a width of approximately 35 km (Cheney and Doblar 1982). The Almeria–Oran Front, as it is known, is thought to be a permanent feature, although its position and intensity are controlled by the degree of development of the Eastern Alboran Gyre (Tintoré et al. 1988). To the east of the Alboran Sea, MAW is concentrated along the northern coast of Africa in the Algerian Current. To its north, northward branches of the MAW form part of various larger-scale cyclonic gyres (Figure 2.3), while smaller anticyclonic gyres are found to the south of the Algerian Current. Waters flowing northwards on both sides of Corsica, the western and eastern Corsica currents, join and form the northern cyclonic gyres in the Gulf of Lions, where the Mistral winds in winter initiate a series of processes leading to the formation of Western Mediterranean Deep Water (WMDW) (e.g. MEDOC Group 1970; Gascard 1978; Leaman and Schott 1991; Robinson and

The Marine Environment: Present and Past

41

39°N

SPAIN

Almeria

Almeria-Oran Front

Gibraltar West Alboran Gyre

East Alboran Gyre

Al

37°N

Algerian Basin

r ge

ian

Current

Algiers

Oran

ALGERIA 35°N 6°W

4°W

2°W

0°

2°E

Fig. 2.5. Schematic illustration of surface circulation in the Alboran Sea (modified from Tintoré et al. 1988).

Golnaraghi 1994; Rohling et al. 1998b; and references therein). MAW enters the eastern Mediterranean through the Strait of Sicily with salinities between 37.0 and 38.5 psu (38.5 is the salinity at the flow reversal boundary; Garzoli and Maillard 1979). It feeds the Ionian Current and Mid-Mediterranean Jet (MMJ) through the Ionian Sea and Levantine basin, respectively. The MMJ bifurcates several times to form a series of cyclonic and

anticyclonic gyres interconnected by jets flowing at speeds of 20–30 cm s−1 (Robinson et al. 1992) (Figures 2.3 and 2.6). One branch of the MidMediterranean Jet flows to Cyprus and then north- and westwards to become the Asia Minor Current (Figure 2.3). It must be noted that salinity values of MAW increase steadily as it travels from west to east, due to continued evaporation (Wüst 1961; Malanotte-Rizzoli and Hecht 1988).

Fig. 2.6. Schematic illustration of the main gyres associated with Atlantic surface flow. IAS = Ionian–Atlantic Stream, CC = Cilician Current, AMC = Asia Minor Current, MMJ = Mid-Mediterranean Jet (modified from Robinson et al. 1992).

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Eelco Rohling et al.

Intermediate-water Circulation

38.5

39.5

500

1000

Depth (m)

During winter, surface waters in the Levantine basin experience enhanced mixing and evaporation as a consequence of strong winds associated with cold, dry air masses tracking the eastern Mediterranean at this time of year (Ozsoy 1981), especially in the Cyprus– Rhodes area. A subsequent combination of low temperatures (15–16◦ C) and high salinities (∼39.2 psu, with extremes to 39.5 psu; Wüst 1960) in surface waters creates favourable conditions for vertical convection and the consequent formation of Levantine Intermediate Water (LIW). This watermass is characterized by a salinity maximum, and spreads throughout the eastern and western Mediterranean basins, forming the major constituent of the Mediterranean Intermediate Water (MIW) (Figure 2.2). It resides between 150 and 600 m water depth, and its transition to surface MAW is marked by a distinct salinity gradient, or halocline (Figures 2.2 and 2.7). There is no comparable source region for intermediate water formation in the western Mediterranean basin. From its source area, LIW/MIW flows westwards, penetrating the Ionian and Adriatic Seas. On approaching the Strait of Sicily, part of the subsurface watermass is recirculated within the eastern basin, while the remainder continues to enter the western Mediterranean basin. The actual ratio of recirculation to efflux remains to be established (Robinson et al. 1992). At the Strait of Sicily, MIW remains distinctive within the water column, although with somewhat reduced temperature (14◦ C) and salinity (38.75%) values compared to those in the LIW source area, due to later admixtures (Garzoli and Maillard 1979). On leaving the Strait, MIW settles between about 200 and 600 m, and splits into main branches that go: into the Tyrrhenian basin; along the western side of Sardinia; and along the AlgerianMoroccan coastlines to exit through the Strait of Gibraltar from the Mediterranean into the North Atlantic. MIW enters the Alboran Sea at a depth between 200 and 600 m, with temperatures and salinities of 13.2◦ C and 38.5% respectively, flowing in a westward direction towards the Strait of Gibraltar at velocities of 1–2 cm s−1 (Parrilla et al. 1986; Richez and Gascard 1986). Since the subsurface outflow through the Strait of Gibraltar displays temperature and salinity values of about 13◦ C and 38.2–38.4 psu, compared to values of 15–16◦ C and 36.1–36.2 psu in the surface (AW) inflow (e.g. Wüst 1960, 1961; Gascard and Richez 1985; Bryden et al. 1994), it is obvious that the Mediterranean experiences both net evaporation and net cooling (Garrett 1994). The subsurface Mediterranean Outflow has

Salinity (p.s.u.) 37.5 0

1500

2000

2500

3000

Fig. 2.7. Typical salinity profiles for the western (dots) and eastern (squares) Mediterranean basins (modified from Rohling and Bryden 1992).

a flux in the order of 1 Sv (Bryden and Kinder 1991) to 1.5 Sv (Béthoux and Gentili 1994) (1 Sverdrup = 1 × 106 m3 s−1 ). It settles between 1,000 and 1,500 m depth in the North Atlantic Ocean (e.g. Wüst 1960; Stommel et al. 1973; Reid 1979; Price et al. 1993).

Deep water Circulation The western and eastern Mediterranean basins each have their own source of deep water, which settles below the MIW. Western Mediterranean Deep Water (WMDW) is formed in the north-west Mediterranean, particularly the Gulf of Lions, and Eastern Mediterranean Deep Water (EMDW) in the Adriatic and Aegean Seas. Today, there is such consistent deep water ventilation from these regions that both the western and eastern Mediterranean are characterized by well-oxygenated deep and bottom waters, with oxygen concentrations

The Marine Environment: Present and Past

typically varying in reported ranges of 4.0–4.7 ml l−1 or 180–210 Ï mol kg−1 (Wüst 1960; McGill 1961; Miller et al. 1970; Schlitzer et al. 1991; Klein et al. 1999; Roether and Well 2001). The following two sections discuss the mechanisms for WMDW and EMDW formation in more detail.

Western Mediterranean Deep Water (WMDW) The Gulf of Lions is the key area for the western Mediterranean deep circulation. The surface circulation in this area is characterized by a distinct cyclonic gyre (MEDOC-group 1970) (Figure 2.3). In winter (January/ February), cold and relatively dry Mistral winds over this region initiate WMDW formation. Three phases can be distinguished: (1) the preconditioning phase, (2) the violent mixing phase; and (3) the sinking and spreading phase (Figure 2.8) (MEDOC Group 1970). During the preconditioning phase, a reduction occurs in the stability of the water column due to winter cooling that leaves surface waters with low temperatures (10–12◦ C), high salinities (38.40 psu), and consequently elevated densities (Wüst 1961; MEDOC Group 1970; Leaman and Schott 1991). At this time mixing occurs in the surface waters but the vertical profile still remains a three-layered one: (1) a relatively fresh and cold surface layer, (2) a warm saline intermediate layer, and (3) a cold and medium-saline deep layer (Figure 2.8). The onset of strong north-westerly Mistral winds (MEDOC-group 1970) initiates an intensification of the basin’s cyclonic circulation, which causes a shallowing of the pycnocline from a usual depth of approximately 200–250 m (Perkins and Pistek 1990) to 2000 m), developing within the centre of the gyre (MEDOC-group 1970; Leaman and Schott 1991). Incidentally, the existence of discrete ‘chimneys’ of deep convective mixing was observed for the first time in this area during the MEDOC study, and similar features have since been recognized in other areas of deep water formation (notably the Norwegian Sea). The geographical extent of the region of deep water formation is characterized at the surface by high salinities (38.4 psu) mixed up from below, from the intermediate water (MEDOC-group 1970).

Preconditioning phase

43

E E

Surface water

Mediterranean intermediate water Western Mediterranean cool, deep water Violent mixing and deep convection phases

E

Surface water

Mediterranean intermediate water Western Mediterranean cool, deep water Fig. 2.8. Schematic illustration of the preconditioning phase, and the violent mixing and deep convection phase. E = Evaporation (modified from Rohling et al. 1998b).

Then follows a phase of sinking and spreading. As the stormy period ceases, the mixed water sinks rapidly to form WMDW (Figure 2.8). The newly formed watermass is characterized by a relatively high oxygen content (4.4–4.7 ml l−1 ), and spreads horizontally between 1,500 and 3,000 m into the Balearic basin and Tyrrhenian Sea (Wüst 1961). On entering the Alboran Sea, WMDW forms a narrow (∼20 km) boundary current flowing westward along the Moroccan coast before entering the Strait of Gibraltar. In the Alboran Sea WMDW reaches speeds of approximately 5 cm s−1 , and it contributes an estimated 0.3 Sv (25%) to the outflow over the Gibraltar Sill (Parrilla et al. 1986; Richez and Gascard 1986).

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Eastern Mediterranean Deep Water (EMDW) Throughout the period of oceanographic observation, until the late 1980s/early 1990s, the Adriatic Sea was found to be the main source area of EMDW formation (Pollak 1951; Wüst 1961; Malanotte-Rizzoli and Hecht 1988; Robinson et al. 1992). In winter, cold and dry north-easterly winds (Bora) cause intense cooling of the North Adriatic shelf waters (Ozsoy 1981), which are of relatively low salinity due to dilution with fresh water from the Po river. The resultant cold waters flow towards the deep south Adriatic basin, where mixing occurs with the warmer but more saline MIW that penetrates the South Adriatic across the Otranto Sill. The mixing of the cold and relatively low-salinity shelf waters with warm and highly saline MIW results in the formation of Adriatic Deep Water (ADW). Although ADW has a lower salinity (300 core-top samples (modified from Hayes et al. 2005).

from Pliocene times. Sapropels range from a few millimetres to more than a metre in thickness, and have been deposited intermittently throughout the Neogene and Quaternary (among countless others: Kullenberg 1952; Olausson 1961; van Straaten

1972; Cita 1973; Cita et al. 1977; Vergnaud-Grazzini et al. 1977; Thunell et al. 1977, 1983; Stanley 1978; Williams et al. 1978; Cita and Grignani 1982; Rossignol-Strick et al. 1982; Rossignol-Strick 1983; Vergnaud-Grazzini 1985; Rohling and Gieskes 1989;

3

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8 5

7

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12

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13

11

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Fig. 2.11. Scanning electron microscope images of the carbonate shells of several planktonic foraminiferal species that live in the Mediterranean Sea. 1, 2 Globigerina bulloides d’Orbigny; 1 umbilical view and 2 spiral view. 3, 4 Globoturborotalita rubescens Hofker; 3 umbilical view and 4 spiral view. 5, 6 Turborotalita quinqueloba (Natland); 5 umbilical view and 6 spiral view. 7 umbilical view and 8 umbilical view. 9, 10 Globigerinoides sacculifer (Brady); 9 umbilical view and 10 spiral view. 11, 12 Globigerinella digitata (Brady); 11 umbilical view and 12 spiral view. 13–15 Globigerinella siphonifera (d’Orbigny); 13 umbilical view, 14 peripheral view, and 15 spiral view. Each scale bar represents 100 Ï m.

The Marine Environment: Present and Past

51

Fig. 2.12. Example of a laminated sapropel in a freshly opened sediment core. The thick dark bed recovered over two core sections represents sapropel S5 from the previous interglacial maximum, 124–119 thousand years ago. The core was recovered during cruise M53-1 of RV Meteor in November–December 2001 (chief scientist Prof. Ch. Hemleben).

Emeis et al. 1991, 1998, 2003; Hilgen 1991a, b; Hilgen et al. 1993, 1995; Rohling 1994; Van Os et al. 1994; Lourens et al. 1992, 1996, 2001; Nijenhuis et al. 1996; Rohling 1999b; Meyers and Negri 2003; and contributions and references therein). Sapropels are commonly marked by an absence of benthic foraminifera, and are preceded by a short interval containing benthic faunas indicative of severe bottom-water oxygen depletion (such faunas sometimes return within, or persist into/through, the sapropel) (e.g. van Straaten 1972; Nolet and Corliss 1990; Verhallen 1991; Rohling et al. 1993b 1997; Nijenhuis et al. 1996; Jorissen 1999; Mercone et al. 2001; Casford et al. 2003; Schmiedl et al. 2003). In marine cores, sapropels are recognizable as beds ranging in colour from dark grey to olive green and black. Exposed in land-sections, sapropels appear in notably darker shades of grey than surrounding beige to blue clays, and

commonly weather into distinct reddish-brown hues. Sapropels may display remarkably well-preserved lamination (Figure 2.12). The shallowest reported occurrence of the four youngest (most cored) sapropels in the open eastern Mediterranean is ∼300 m (Shaw and Evans 1984; Rohling and Gieskes 1989; Rohling et al. 1993a ). In the Adriatic Sea, the upper depth limit seems to have been at a deeper level, below 400 m (Jorissen et al. 1993). In the Aegean Sea, sapropels are found up to 120 m water depth (Perissoratis and Piper 1992; Casford et al. 2002). Following almost six decades of research on Mediterranean sapropels since their initial discovery in marine sediment cores recovered during the Swedish Deep-Sea Expedition of 1946–7, a general (but not unanimous) consensus has emerged that sapropels were formed during times with a combination of (1) enhanced abundances of organic matter sinking

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and Grignani 1982; Rohling 1994; Emeis et al. 1998; Cramp and O’Sullivan 1999; Rohling et al. 2004). These impacts will be discussed below, and are summarized in Figure 2.13 (modified after Rohling 1994).

from surface waters (i.e. export production), and (2) reduced deep water ventilation due to diminished excess evaporation from the Mediterranean basin caused by enhanced freshwater discharge (for overviews, see Cita

45°N

40°

A B

35°

C

30°N 5°W

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Present day (pre-1988)

S

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S

Aegean Sea

N

Depth (m)

0

AeDW

ADW

700

A W

B Open Eastern Mediterranean

E

MIW

AeDW ADW

C Fig. 2.13. Schematic presentation of the changes in subsurface circulation patterns between the present day and times of sapropel formation. The three profiles presented summarize information obtained from analytical and modelling studies from north to south through the Adriatic and Aegean basins, and from west (Strait of Sicily) to east (near Cyprus) through the open eastern Mediterranean. MIW stands for Mediterranean Intermediate Water; ADW for Adriatic Deep Water; AeDW for Aegean Deep Water; AIW for Adriatic Intermediate Water; AeIW for Aegean Intermediate Water; ODW for Old (isolated) Deep Water. Modified from Rohling (1994) and Myers et al. (1998).

The Marine Environment: Present and Past

S

Adriatic Sea

Sapropel-time S N

Aegean Sea

53

N

0

Depth (m)

AeIW AIW

ODW

700

ODW

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Open Eastern Mediterranean AeIW AIW

ODW C Fig. 2.13. (cont.)

Evidence for elevated export production has been compiled through a combination of proxy data that includes phyto- and zoo-plankton abundances, stable isotope gradients, organic-carbon accumulation and composition, and Ba/Al ratios in the sediment (among many others: Cita and Grignani 1982; Rohling and Gieskes 1989; Castradori 1993; Higgs et al. 1994; Thomson et al. 1995, 1999; Van Os et al. 1994; Kemp et al. 1999; Mercone et al. 2000, 2001; Rohling et al. 2004). Productivity increases during the deposition of sapropels appear to have been of a (temporally integrated) basin-wide nature, especially in the form of a Deep Chlorophyll Maximum (Rohling and Gieskes 1989; Castradori 1993; Rohling 1994; Kemp et al. 1999; Corselli et al. 2002), although on shorter time-scales there may have been considerable spatial ‘patchiness’ (Casford et al. 2003). The development of a Deep Chlorophyll Maximum with high export production at sapropel times has been ascribed to hydrographic rearrangements in response to the decrease in buoyancy loss from the Mediterranean at times of enhanced freshwater input. Notably, the reduced surface buoyancy loss is thought to have caused a shoaling of the surface–intermediate water interface from its present

depth below the zone of light penetration (euphotic zone) to a depth within the euphotic layer (Rohling and Gieskes 1989; Rohling 1991b, 1994; Myers et al. 1998). This would allow nutrients stored within the subsurface waters to become utilized for production at the base of the euphotic layer. There is a continuing debate about the ultimate supply of the nutrients that could have supported extensive organic carbon burial in the sediments. Early work concentrated on riverine nutrient input at times of sapropel deposition, but biogeochemical modelling suggests that river-input would be insufficient if the nutrient budget were at steady state during sapropel formation (Stratford et al. 2000). However, recent work has suggested that the basin may have accumulated nutrients over as much as 1,500 years prior to the onset of organic carbon burial, so that the nutrient budget during sapropel deposition ought to be considered as a product of accumulation over much longer timescales, and so was not at steady state (Casford et al. 2002). Strong evidence for enhanced freshwater influx into the eastern Mediterranean at sapropel times comes from negative anomalies in stable oxygen isotope ratios measured on the calcium-carbonate shells of

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+2.0

0

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Sapropel geochronology

Oxygen Isotope Record G. ruber

Age (Ma)

Sapropels

Depth (m)

Core RC9-181

Precession

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+ min

max

S1 1 2 3 4 5 6

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planktonic foraminifera that live in near-surface habitats. Fresh water has distinctly low oxygen isotope ratios compared with sea water, and especially the fresh waters derived from heavy (monsoon-type) rainfalls that are isotopically very light. Freshwater floods to the Mediterranean surface waters therefore cause light isotope anomalies in the surface-dwelling foraminifera (e.g. Vergnaud-Grazzini et al. 1977; Thunell and Williams 1983, 1989; Jenkins and Williams 1984; Ganssen and Troelstra 1987; Kallel et al. 1997a, b; Tang and Stott 1993; Rohling and De Rijk 1999a, b; Emeis et al. 1998, 2000, 2003; Rohling et al. 2004). Sedimentary Ti/Al ratios, palaeomagnetic data, and clay mineralogical studies confirm that times of sapropel deposition were characterized by humid climates with high runoff, whereas intervening times were arid with reduced riverine and enhanced wind-blown sediment supply (e.g. Krom et al. 1999; Foucault and Mélières 2000; Wehausen and Brumsack 2000; Lourens et al. 2001; Larrasoaña et al. 2003). A particularly important discovery concerned the temporal coincidence between sapropel occurrences and insolation-driven monsoon maxima, affecting the eastern Mediterranean via changes in Nile discharge (Rossignol-Strick et al. 1982; Rossignol-Strick 1983, 1985). These authors approached the problem by specifying an index for monsoon intensity (‘monsoon index, M’) as a function of two parameters, namely insolation

Fig. 2.14. Phase relationships between the sapropel record and associated ‰18 O record from core RC9-181 and the orbital cycles of precession and eccentricity (modified from Hilgen 1991a).

at the (north) Tropic of Cancer (IT ), and the insolation difference between the Tropic of Cancer and the equator (IT –IE ), so that M = 2IT –IE . The variation in the index value was considered relative to the value of AD 1950 (Rossignol-Strick 1985). This pioneering work instigated an intensive search into the timing of sapropel formation over their full temporal range, which confirmed that sapropels were always formed at times when perihelion falls in boreal summer (‘precession minima’, relative to ‘maxima’ that represent the present configuration with perihelion in boreal winter). It was also observed that not all precession minima have sapropels, but that they instead occur in discrete clusters. Each cluster was found to represent times of maximum orbital eccentricity, in agreement with eccentricity modulation of the impact of precession (Hilgen 1991a , b; Hilgen et al. 1993, 1995; Lourens et al. 1996, 2001) (Figure 2.14). Numerical climate modelling corroborated the impact of precession and eccentricity on monsoon intensity (e.g. Kutzbach 1985; Kutzbach and Guetter 1986; COHMAP 1988). In essence, insolation changes affect the monsoons: air rises over a hot surface, giving low surface pressure, while it descends over a cool surface, giving high surface pressure. Because land has much lower thermal inertia than ocean, land surfaces experience a much stronger annual fluctuation in both temperature and pressure

The Marine Environment: Present and Past

than ocean surfaces. During periods with enhanced seasonal insolation contrasts, the higher summer insolation increases surface temperatures especially over land, and this in turn amplifies the atmospheric pressure differences between land and sea. In addition to this direct radiative forcing, the preceding winter conditions also play a role, due to the thermal inertia of the ocean. The slow response of oceanic temperatures on seasonal timescales amplifies the land–sea temperature contrast from direct solar heating, and thus enhances the land– sea pressure differences. In summer, the strong land (low) to sea (higher) pressure difference leads to surface air flow from ocean to land. This air flow is moistureladen, because of evaporation over the ocean. The air expands and cools as it rises over the land, a process that is accelerated if the air masses are forced up by mountain ranges. The cooling causes the air masses to shed their vapour as rain. Condensation releases heat, which amplifies the process by enhancing the ascending motion in the air column. Thus, a zone develops of highfrequency and high-intensity monsoonal rains. We emphasize that the above description of the summer monsoon in terms of surface thermal forcing (i.e. as a super sea breeze) represents a strongly simplified generalization. In reality, the low-pressure cell over land derives much of its intensity and continuity from dynamical effects related to the mean high-level wind flow in the atmosphere (at the 500 millibar level, or approximately at 5.5 km height), especially in the case of the Indian/Asian monsoon. As an extra complication, it is thought that the strength of the trade winds in the opposite (winter) hemisphere may determine a ‘push’ across the equator into the summer monsoonal low. Despite its schematic nature, however, the thermal concept offers a reasonable representation of the general features of the African monsoon. Over Africa, the axis of low pressure at the surface (‘the monsoonal low-pressure trough’) follows the seasonal march of the sun at its high point (zenith), which reaches the Tropic of Cancer at the summer solstice. This seasonal swing over the band of monsoon-influenced latitudes in Africa can be smooth because most of Sahelian and Saharan North Africa is relatively flat. The influence of ‘push’ effects by the Southern (winter) Hemisphere trade winds on the North African summer monsoon was accounted for in the monsoon intensity calculations of Rossignol-Strick and co-workers by inclusion of an austral winter insolation gradient (GS ) between 20 and 70◦ S, so that GS = I20 –I70 (Rossignol-Strick 1985). Within the context of monsoon intensification, it is relevant to briefly review reconstructions of the Nile and the Sahara since the LGM, as summarized by Adamson

55

et al. (1980) and Williams et al. (2000). These authors report that, from the LGM until roughly 12,500 years BP , Nile discharge was very low. The White Nile was a seasonal, intermittent river until ∼12,500 years BP, when Lake Victoria overflow developed and the ‘buffering’ Sudd swamps in Sudan became established, which ensured a more regular, perennial discharge from the White Nile. From ∼12,500 years BP, there was an (intermittent) period with very high discharge, associated with the early–mid Holocene monsoon maximum that developed during the insolation maximum of that time. This maximum ended with a development towards generally much drier conditions around 5,000 years BP, heralding the development of the modern Nile regime. These trends are supported by general findings that astronomical forcing affects not only the intensity of the African monsoon, but also its spatial influence, causing strong reductions in the size of the Sahara desert by northward migration of its southern margin. The ‘greening’ of the Sahara is a well-known response in numerical climate models that include vegetation– climate feedback mechanisms (Brovkin et al. 1998; Claussen et al. 1998). The concept is supported by a wide variety of field observations: rock-art and animal, human, and vegetation remains in the central Sahara; a massive expansion of Lake Chad; the presence of substantial palaeolakes in currently hyperarid areas such as the Oyo depression of NW Sudan; and activation of large-scale systems of presently inactive wadis (e.g. Pachur and Braun 1980; Gaven et al. 1981; Ritchie et al. 1985; McKenzie 1993; Szabo et al. 1995; PetitMaire and Guo 1997; Pachur 2001; Gasse 2000, Hoelzmann et al. 2000; Williams et al. 2000; Mandel and Simmons 2001; Hassan 2002; and many references therein). A recent study suggested on the basis of stable oxygen isotope data that the monsoon front penetrated sufficiently far northwards during the insolation maximum of the previous interglacial maximum to have caused significant precipitation to the north of the central Saharan watershed (∼21 ◦ N) (Rohling et al. 2002b). In that case, significant runoff would have reached not only the eastern Mediterranean via the Nile River, but also the wider North African margin. A study concerned with aeolian dust variations over the last 3 million years supported that scenario for all substantial insolation maxima (Larrasoaña et al. 2003). Archaeological observations around exclusively rain-fed depressions on the Libyan Plateau suggest that monsoonal summer rains of central Africa periodically penetrated at least as far north as Kharga (roughly 25 ◦ N) during the early–mid Holocene, despite the fact that conditions during that pluvial phase seem to have remained

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drier than during earlier Quaternary pluvial phases (Mandel and Simmons 2001). The observation of Mandel and Simmons (2001) that the Holocene monsoon maximum was of a relatively low intensity compared with previous Quaternary monsoon maxima has been corroborated by recent work to quantify the Holocene and previous interglacial (Eemian) monsoon impacts on the freshwater budget of the eastern Mediterranean (Rohling 1999b; Rohling et al. 2004). Eastern Mediterranean surface-water oxygen isotope (‰18 O) data show two very distinct ‘peaks’ in the Eemian monsoon maximum, separated by an ‘interruption’ that lasted about 800 years. A mixed-layer ‰18 O box model to quantify freshwater flooding during the two Eemian monsoon peaks suggests that the basinaveraged Mediterranean excess of evaporation over freshwater input was reduced to 5–45 per cent (older peak) and 35–60 per cent (younger peak), relative to the present (Rohling et al. 2004). It also suggests that the interruption between the two peaks was characterized by excess evaporation at levels very close to present-day values. Using a similar technique, the excess evaporation during the early–mid Holocene monsoon maximum was estimated at about 65 per cent of the present value (Rohling 1999b). Using a different technique, based on an ocean general circulation model (OGCM), Myers (2002) suggests that the excess evaporation value for the Holocene monsoon maximum ranged below 80 per cent and most likely around 20–40 per cent of the present-day value. As yet, eastern Mediterranean ‰18 O records through the Holocene monsoon maximum have revealed only weak indications of a monsoon interruption, but neodymium (Nd) isotopes seem more conclusive that such an interruption did occur during the Holocene (Scrivner et al. 2004). African lake levels also clearly demonstrate an arid interlude, dated between about 8.5 and 7.8 ka BP (Gasse 2000), coincident with a cooling event of ∼500-year duration over the Aegean and Adriatic seas (see e.g. Rohling et al. 2002a ). There is a growing body of evidence showing that it was not just the monsoon system that was intensified at times of sapropel deposition. Records of pollen and spore abundances from terrestrial vegetation suggest high abundances of species that require wet summers around the Northern Borderlands of the Eastern Mediterranean (NBEM) at times of sapropel formation (Rossignol-Strick 1987, 1995; Wijmstra et al. 1990; Rohling and Hilgen 1991; Tzedakis 1993; Mommersteeg et al. 1995; Frogley et al. 1999). Winter precipitation is thought to have been increased in the NBEM as well (Wijmstra et al. 1990). Isotope studies on speleothems corroborate

the inferred increase in precipitation (Bard et al. 2002; Matthews et al. 2000; Bar-Matthews et al. 1999, 2000, 2003), as do elevated lake levels (e.g. Digerfeldt et al. 2000). The conditions at times of sapropel formation would therefore appear to have been considerably different from the typical dry-summer climate that characterizes the area today. Direct precipitation from the African and Indian monsoon systems is unlikely to have penetrated into the Mediterranean basin, demonstrating that the moisture for precipitation in the NBEM derived from regional processes, probably in the form of Mediterranean depressions (Rohling and Hilgen 1991). This notion was corroborated by isotopic characteristics of speleothems in Soreq Cave, Israel, which demonstrate a local Mediterranean moisture origin (Matthews et al. 2000; Bar-Matthews et al. 2003). The summer flux of Mediterranean moisture at times of insolation maxima affected even the northernmost tip of the Red Sea, where it has been described as the ‘Mediterranean monsoon’ (Arz et al. 2003). Importantly, this process would not likely have affected the Mediterranean Sea’s overall hydrological budget very significantly—although any precipitation would have led to runoff into the basin, the original evaporative loss took place from the same basin. Any transport into another basin’s watershed area (e.g. the Red Sea, Jordan Valley and Dead Sea, or Tigris/Euphrates and Persian Gulf) would imply that the regional ‘humidity’ in the NBEM might even have coincided with a slight increase in net evaporative loss from the Mediterranean. Importantly, however, the process would reflect considerable freshwater redistribution in a generally eastward direction within the Mediterranean basin, so that the hydrological budget may have been substantially affected on local scales and in terms of regional gradients. The (especially monsoon-related) reduction of Mediterranean excess evaporation would have caused a reduction in the salinity of newly formed intermediate water. Numerical modelling suggests that intermediatewater formation is likely to have shifted from a normal salt-dominated LIW mode, to a temperature-dominated mode driven from the Adriatic Sea at times of sapropel formation (Myers et al. 1998; Myers 2002). Stable isotope data for planktonic foraminiferal species with different depth habitats from the Aegean Sea corroborate that notion (Casford et al. 2002). The collapse of the first, salt-driven, stage of the deep-ventilation ‘motor’ would have caused any new deep water to form at much lower salinities (hence, lower densities) than it did in times before the monsoon intensification (Rohling 1994; Myers et al. 1998; Myers 2002). Thus,

The Marine Environment: Present and Past

newly formed deep water masses could not displace the existing, denser, ‘old’ deep waters (ODW) formed before the freshwater flooding (Figure 2.13). As a result, the ODW became poorly ventilated, and eventually oxygen depleted due to continuing remineralization of sinking organic matter (for an overview, see Rohling 1994). At least down to ∼2000 m depth some occasional ventilation may have persisted during the deposition of several sapropels, and down to those depths the occurrence of truly anoxic conditions was probably restricted to spatially discontinuous ‘blankets’ over the sea-floor topography (Casford et al. 2003).

Centennial- to Millennial-scale Variability As mentioned previously, deep ventilation in the Mediterranean basin is strongly affected by wintertime, orographically channelled, northerly outbursts of cold polar and continental air over the northern sectors of the basin. Fluctuations in the intensity and frequency of such events are also reflected in temperature proxy data. Terrestrial and marine Mediterranean palaeoclimate and palaeoceanographic proxy records that are resolved on centennial timescales have been found to reflect multi-centennial to millennial fluctuations (Rohling et al. 1998b, 2002a, b, 2004; Paterne et al. 1999; Cacho et al. 1999, 2000, 2001, 2002; Allen et al. 1999; Combourieu-Nebout et al. 2002; SánchezGoñi et al. 2002; Tzedakis et al. 2004). These have often been related to climatic oscillations in the wider North Atlantic realm by correlation with the ice-‰18 O records and non-sea-salt ion series (dust) from the welldated GISP2 and GRIP ice cores (Greenland summit). Particularly strong cooling events have been described for the northern sectors of the Mediterranean at times coincident with the North Atlantic ‘Heinrich Events’ of massive ice-rafting (Rohling et al. 1998b; Paterne et al. 1999; Cacho et al. 1999, 2000, 2001, 2002). Within the last glacial cycle, periods of intensified or more frequent northerly cooling events in the Mediterranean generally correlate with the DansgaardOeschger (DO) stadials (cold episodes), and the most intense events correlate with the most intense DO stadials, which were marked in the North Atlantic by the Heinrich Events. Highly resolved records from the Mediterranean region furthermore indicate that the cool periods were characterized by enhanced aridity (e.g. Allen et al. 1999; Sánchez-Goñi et al. 2002; Combourieu-Nebout et al. 2002; Tzedakis et al. 2004; Hoogakker et al. 2004).

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The most recent sapropel (S1) formed between about 9,000 and 6,000 calibrated years BP, in association with the monsoon maximum of the current interglacial period. Detailed work has found an ‘interruption’ in the deposition of the S1 sapropel, marking a period of improved deep water ventilation that spans several centuries around 8,500–8,000 years BP (van Straaten 1966, 1970, 1972, 1985; Rohling et al. 1997; De Rijk et al. 1999; Geraga et al. 2000; Mercone et al. 2000, 2001; Casford et al. 2001, 2002, 2003; Rohling et al. 2002a ). This ventilation event coincided with intensified cooling over the Adriatic and Aegean seas, which has been related to an increase in the intensity or frequency of northerly cold outbursts over those regions, in association with a widespread North Atlantic cold event (Rohling et al. 2002a ). Similarly, the ending of sapropel deposition was marked by a cooling, around 6,000 years BP (Rohling et al. 1997; Geraga et al. 2000; Casford et al. 2001, 2002, 2003; Mercone et al. 2001; Rohling et al. 2002a). Pollen data confirm that the cooling periods were normally marked by enhanced aridity (e.g. Rossignol-Strick 1995; Geraga et al. 2000). Episodically improved deep water ventilation within times of generally poor ventilation (sapropel conditions) has since been inferred for a large number of sapropels, suggesting that climatic variability on short timescales persisted even in these periods of generally warm and humid conditions in the Mediterranean region (Casford et al. 2003). The above might give the impression that sapropels resulted from monsoon maxima extending over several millennia, while some internal variability occurred due to intermittent cooling events originating from the north, in association with polar or temperate climate events. This would be misleading, because the centennial-scale episodes of increased cooling from the north are well known to have been associated with severe reductions in monsoon runoff. One likely example is the 800-year monsoon interruption within Eemian sapropel S5 (e.g. Cane et al. 2002; Rohling et al. 2002b, 2004). Monsoon-fed African lakes show a similar interruption within the Holocene monsoon maximum, as part of a series of distinct and abrupt periods of low lake levels that coincide in time with the northerly coolings recorded in the Mediterranean, as described above. The particularly pronounced period of low lake levels between 8.5 and 7.8 ka BP (Gasse 2000) coincides closely with the interruption of sapropel S1, and recent compilations have found this to be a widespread interval of climate deterioration throughout (at least) the Northern Hemisphere (Mayewski et al. 2004; Rohling and Pälike 2005). Furthermore, Egyptian archaeological

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records for the Holocene indicate dramatic turnovers related to strong fluctuations in intensity and frequency of Nile flooding (Hassan 1997, 2002, and references therein). Clearly, some fundamental, if elusive, connection exists between cooling from the north and reductions in (African) monsoon intensity. Hence, it is worthwhile to spend some time looking at variability in the wind-blown dust flux into the Mediterranean, as a measure of North African climate variability. At present, the northward transport of aeolian (windblown) dust over the Mediterranean is linked to the presence of cyclones over the basin (Moulin et al. 1997), and most Saharan dust deposition over southern Europe occurs with precipitation (Bücher and Lucas 1984; Bergametti et al. 1989a; Loye-Pilot et al. 1989; Guerzoni et al. 1992). Important source areas of dust transport to Western Europe are Algeria, the Western Sahara, and the Moroccan Atlas (Molinaroli 1996; Avila et al. 1997; Goudie and Middleton 2001; Chapter 14). Weldeab et al. (2002) use Si/Al and Ti/Al ratios as well as Sr and Nd isotopes to show that the Saharan terrigenous input into the western Mediterranean Sea is predominantly from the southwest (Morocco/north-western Algeria) and south-east (Tunisia/western Libya) during interglacial periods and from the southern Saharan/Sahelian region during glacial times. Eastern Libya and Egypt are the important source areas of aeolian dust to the eastern Mediterranean basin. Overall, terrigenous input into the Mediterranean at glacial times greatly exceeded that of interglacial times (Weldeab et al. 2002). Fluvial sediment yields were also higher (Macklin et al. 2002). A continuous 3 million-year record of dust supply from the northern Sahara into the eastern Mediterranean, developed from sediment magnetic data, consistently shows dust-flux minima at times of Northern Hemisphere insolation/monsoon maxima (Larrasoaña et al. 2003). These minima were related to northward penetration of the African summer monsoon front beyond the central Saharan watershed (∼21 ◦ N), as proposed previously on the basis of Mediterranean oxygen isotope data (Rohling et al. 2002b). Such northward penetration of the African summer monsoon agrees with a broad expansion of (savannah-like) vegetation cover shown in both observations and modelling experiments (‘greening of the Sahara’: Claussen et al. 1998; Brovkin et al. 1998; Irizarry-Ortiz et al. 2003; and references therein). Consequently increased soil cohesiveness throughout large areas of the northern Sahara would cause a decrease in dust production, similar to modern conditions in the Sahel (Middleton 1985).

A high-resolution record of lithogenic fraction variability from the Alboran Sea has revealed millennialto submillennial-scale oscillations. These correlate with Atlantic Dansgaard-Oeschger stadials and Heinrich Events, and were characterized by increases in the northward transport of Saharan dust (Moreno et al. 2002). Similar increases in Saharan dust supply at times of DO stadials have been inferred from sediment cores throughout the wider western Mediterranean (Hoogakker et al. 2004), and detailed magnetic susceptibility records for cores taken in the eastern Mediterranean (pp. 3/28–3/31 in Hemleben et al. 2003) suggest that a millennial-scale aeolian dust signal may also be preserved in that basin. At this stage, we do not infer that the monsoon penetration model inferred for the longer-term (Milankovitch-scale) dust cycles should apply to the shorter-term (sub-Milankovitch) events. Instead, the shorter-term dust-flux variations may well be controlled by cyclone activity, as it appears to be on interannual to decadal time scales (Moulin et al. 1997). An important control on cyclogenesis within the Mediterranean basin is exerted by cold (arid) air outbursts over the northern sectors (Chapter 3), and this may be the mechanism underlying the correlation between dust cycles in the Mediterranean and DO events in the North Atlantic region. This interpretation remains speculative, however, until more process-oriented research leads to a more detailed understanding.

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S1, the most recent eastern Mediterranean sapropel. Geochimica et Cosmochimica Acta 59: 3487–501. Mercone, D., de Lange, G. J., and van Santvoort, P. J. M. (1999), Review of recent advances in the interpretation of eastern Mediterranean sapropel S1 from geochemical evidence. Marine Geology 153: 77–89. Thunell, R. C. (1986), Pliocene-Pleistocene climatic changes: evidence from land-based and deep-sea marine records. Mem. Societa Geologica Italiana 31: 135–43. and Williams, D. F. (1983), Paleotemperature and paleosalinity history of the eastern Mediterranean during the late Quaternary. Palaeogeography, Palaeoclimatology, Palaeoecology 44: 23–39. (1989), Glacial-Holocene salinity changes in the Mediterranean Sea: hydrographic and depositional effects. Nature 338: 493–6. and Kennett, J. P. (1977), Late Quaternary paleoclimatology, stratigraphy and sapropel history in eastern Mediterranean deep-sea sediments. Marine Micropaleontology 2: 371–88. and Cita, M. B. (1983), Glacial anoxia in the eastern Mediterranean. Journal of Foraminiferal Research 13: 283–90. and Howell, M. (1987), Atlantic-Mediterranean water exchange during the late Neogene. Paleoceanography 2: 661–78. Rio, D., Sprovieri, R., and Vergnaud-Grazzini, C. (1991), An overview of the post-Messinian paleoenvironmental history of the western Mediterranean. Paleoceanography 6: 143–64. Tintoré, J. D., La Violette, P. E., Blade, I., and Cruzado, A. (1988), A study of an intense density front in the eastern Alboran Sea: The Almeria-Oran Front. Journal of Physical Oceanography 18: 1384–97. Trewartha, G. T. (1966), The Earth’s Problem Climates. Methuen, London. Trigo, I. F., Davies, T. D., and Bigg, G. R. (1999), Objective climatology in the Mediterranean region. Journal of Climate 12: 1685–96. – (2000), Decline in Mediterranean rainfall caused by weakening of Mediterranean cyclones. Geophysical Research Letters 27: 2913–16. Troelstra, S. R. (1987), Late Quaternary sedimentation in Tyro and Kretheus Basins, southeast of Crete. Marine Geology 75: 77–91. Tzedakis, P. C. (1993), Long-term tree populations in northwest Greece through multiple climatic cycles. Nature 364: 437–40. (1999), The last climatic cycle at Kopais, central Greece. Journal of the Geological Society, London 156: 425–34. Frogley, M. R., Lawson, I. T., Preece, R. C., Cacho, I., and de Abreu, L. (2004), Ecological thresholds and patterns of millennial-scale climate variability: the response of vegetation in Greece during the last glacial period. Geology 32: 109–12. UNDP/UNESCO (1978), Coastal Protection Studies. Project Findings and Recommendations. Paris, UNDP/EGY/73/063. Van Os, B. J. H., Lourens, L. J., Hilgen, F. J., De Lange, G. J., and Beaufort, L. (1994), The formation of Pliocene sapropels and carbonate cycles in the Mediterranean: diagenesis, dilution, and productivity. Paleoceanography 9: 601–17. van Straaten, L. M. J. U. (1966), Micro-malacological investigation of cores from the southeastern Adriatic Sea. Proceedings, Koninklijke Nederlandse Akademie Sciences, Serie B 69: 429–45.

(1970), Holocene and late-Pleistocene sedimentation in the Adriatic Sea. Geologische Rundschau 60: 106–31. (1972), Holocene stages of oxygen depletion in deep waters of the Adriatic Sea, in D. J. Stanley (ed.), The Mediterranean Sea. Dowden, Hutchingson, & Ross, Stroudsburg, Pa., 631–43. (1985), Molluscs and sedimentation in the Adriatic Sea during late-Pleistocene and Holocene times. Giornale Geologia, Serie 3a 47: 181–202. Vergnaud-Grazzini, C. (1985), Mediterranean late Cenozoic stable isotope record: stratigraphic and paleoclimatic implications, in D. J. Stanley and F. C. Wezel (eds.), Geological Evolution of the Mediterranean Basin. Springer, New York, 413–51. Ryan, W. B. F., and Cita, M. B. (1977), Stable isotope fractionation, climatic change and episodic stagnation in the eastern Mediterranean during the late Quaternary. Marine Micropaleontology 2: 353–70. Borsetti, A. M., Cati, F., Colantoni, P., d’Onofrio, S., — Saliège, J. F., Sartori, R., and Tampieri, R. (1988), Palaeoceanographic record of the last deglaciation in the Strait of Sicily. Marine Micropaleontology 13: 1–21. Capotondi, L., and Lourens, L. J. (1993), A refined Pliocene to early Pleistocene chronostratigraphic frame at ODP Hole 653A (West Mediterranean), Marine Geology 117: 329–49. Verhallen, P. J. J. M. (1991), Late Pliocene to Early Pleistocene Mediterranean mud-dwelling foraminifera; influence of a changing environment on community structure and evolution. Utrecht Micropaleontological Bulletin 40. Voelker, A. H., et al. (2002), Global distribution of centennialscale records for marine isotope stage (MIS) 3: a database. Quaternary Science Reviews 21: 1185–212. Wahby, S. D. and Bishara, N. F. (1981), The effect of the river Nile on Mediterranean water, before and after the construction of the High Dam at Aswan, in J. M. Martin, J. D. Burton, and D. Eisma (eds.), River Inputs to Ocean Systems. UNEP-UNESCO, Switzerland, 311–18. Wallmann, K., Suess, E., Westbrook, G. H., Winkler, G., Cita, M. B., and the MEDRIFF consortium (1997), Salty brines on the Mediterranean sea floor. Nature 387: 31–2. Wehausen, R. and Brumsack H. J. (2000), Chemical cycles in Pliocene sapropel-bearing and sapropel-barren eastern Mediterranean sediments. Palaeogeography, Palaeoclimatology, Palaeoecology 158: 325–52. Weldeab, S., Emeis, K. C., Hemleben, C., and Siebel, W. (2002), Provenance of lithogenic surface sediments and pathways of riverine suspended matter in the Eastern Mediterranean Sea: evidence from 143 Nd/144 Nd and 87 Sr/86 Sr ratios. Chemical Geology 186: 139–49. Wijmstra, T. A., Young, R., and Witte, H. J. L. (1990), An evaluation of the climatic conditions during the Late Quaternary in northern Greece by means of multivariate analysis of palynological data and comparison with recent phytosociological and climatic data. Geologie en Mijnbouw 69: 243–51. Williams, D. F., Thunell, R. C., and Kennett, J. P. (1978), Periodic fresh-water flooding and stagnation of the eastern Mediterranean during the late Quaternary. Science 201: 252–4. Williams, M. A. J., Adamson, D. A., Cock, B., and McEvedy, R. (2000), Late Quaternary environments in the White Nile region, Sudan. Global and Planetary Change 26: 305–16. Wüst, G. (1960), Die tiefenzirkulation des Mittelländischen Meeres in den Kernschichten des Zwischen-und des Tiefenwassers. Deutsche Hydrographische Zeitschrift 13: 105–31.

The Marine Environment: Present and Past Wüst, G. (1961), On the vertical circulation of the Mediterranean Sea, Journal of Geophysical Research 66: 3261–71. Yüce, H. (1995), Northern Aegean water masses. Estuarine, Coastal and Shelf Science 41: 325–43. Zachariasse, W. J. and Spaak, P. (1983), Middle Miocene to Pliocene paleoenvironmental reconstruction of the Mediterranean and adjacent Atlantic Ocean: planktonic foraminiferal record of southern Italy, in J. E. Meulenkamp (ed.) Reconstruction of marine paleoenvironments. Utrecht Micropaleontological Bulletin 30: 91–110.

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Zijderveld, J. D. A., Langereis, C. G., Hilgen, F. J., and Verhallen, P. J. J. M. (1989), Early Late Pliocene biochronology and surface water temperature variations in the Mediterranean. Marine Micropaleontology 14: 339–55. Gudjonsson, L., Hilgen, F. J., Langereis, C. G., Lourens, L. J., Verhallen, P. J. J. M., and Zijderveld, J. D. A. (1990), Late Gauss to Early Matuyama invasions of Neogloboquadrina atlantica in the Mediterranean and associated record of climatic change. Paleoceanography 5: 239–52.

This chapter should be cited as follows Rohling, E. J., Abu-Zied, R. H., Casford, J. S. L., Hayes, A., and Hoogakker, B. A. A. (2009), The marine environment: present and past, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 33–67.

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3

The Climate System Andrew Harding, Jean Palutikof, and Tom Holt

Introduction The Mediterranean region has a highly distinctive climate due to its position between 30 and 45◦ N to the west of the Euro-Asian landmass. With respect to the global atmospheric system, it lies between subtropical high pressure systems to the south, and westerly wind belts to the north. In winter, as these systems move equatorward, the Mediterranean basin lies under the influence of, and is exposed to, the westerly wind belt, and the weather is wet and mild. In the summer, as shown in Figure 3.1, the Mediterranean lies under subtropical high pressure systems, and conditions are hot and dry, with an absolute drought that may persist for more than two or three months in drier regions. Climates such as this are relatively rare, and the Mediterranean shares its winter wet/summer dry conditions with locations as distant as central Chile, the southern tip of Cape Province in South Africa, southwest Australia in the Southern Hemisphere, and central California in the Northern Hemisphere. All have in common their mid-latitude position, between subtropical high pressure systems and westerly wind belts. They all lie on the westerly side of continents so that, in winter, when the westerly wind belts dominate over their locations, they are exposed to rain-bearing winds. In the Köppen classification (Köppen 1936), these climates are known as Mediterranean (Type Cs, which is subdivided in turn into maritime Csb and continental Csa). The influence of the Mediterranean Sea means that the Mediterranean-type climate of the region extends much further into the continental landmass than elsewhere, and is not restricted to a narrow ocean-facing strip. Nevertheless, within the Mediterranean region

climate is modified by position and topographic influences can be important. The proximity of the western Mediterranean to the Atlantic Ocean gives its climate a maritime flavour, with higher rainfall and milder temperatures throughout the year. The eastern Mediterranean lies closer to the truly continental influences of central Europe and Asia. Its climate is drier, and temperatures are hotter in summer and colder in winter than in the west. Annual rainfall is typically around 750 mm in Rome, but only around 400 mm in Athens (Figures 3.2 and 3.3a). The southern shore of the Mediterranean Sea is drier and hotter than the northern shore. In the extreme south-east, which lies permanently under the subtropical high pressure belts, the climate becomes arid and hot, and no longer falls into the Köppen Mediterranean class. Here the Sahara desert meets the Mediterranean coast and mean annual rainfall in Alexandria is just 178 mm, with five absolutely dry months each year from May to September (Figures 3.2 and 3.3b). Winter and summer mean temperature and mean monthly precipitation are shown for five locations (including Lisbon on the Atlantic seaboard) in the region in Figure 3.2. Annual rainfall across the Mediterranean region is shown in Figure 3.3a with a general decrease in values moving to the south and east although this is modified by topography. In the high mountains of the Mediterranean the climate can be very humid indeed with annual precipitation exceeding 3,000 mm (Chapter 12). Altitude also plays an important role and climate can change dramatically moving inland from the coast. Snow and freezing conditions are rare in low-lying areas, but in mountainous regions such as the Apennines of Italy and Pindus of Greece winters can be severe,

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January

July

H 1020 H

40°N

40°N

102

0

L

H

20 10

1000

L

MF

H

L

H 20°N

1010

20°N L

ITF

1015 L

1010

ITF 20°W L 0°

H

0° 20°W

1010

20°E

H

40°E

40°E 0°

Monthly precipitation (mm)

20°E

50–400 >400

Fig. 3.1. The location of the Mediterranean region in relation to the large-scale atmospheric circulation (modified from Barry and Chorley 1992). See Figure 2.4.

and thick snow is common, with glaciers present in both the Pyrenees and Apennines (Chapter 12). Indeed, there are places in the Mediterranean where tourists can spend the morning skiing and the afternoon on the beach, including ski resorts in the Troodos Mountains of Cyprus, and Faraya-Mzaar in Lebanon. Under these conditions, the critical climatic factor for the region is rainfall (Figure 3.3). Although people may feel that winters are too cold, and that in summer the heat becomes excessive, in fact temperatures are rarely a limiting factor to activities. But rainfall, or rather lack of it, may be. This becomes a particular problem on the many small inhabited islands of the Mediterranean, many of which are magnets for tourists in the summer, when there is no rain. Conflicts arise between requirements for domestic supply, irrigation, and tourism (Chapter 21). The water supply comes mainly from groundwater, and salinization and mineralization of the supply is common (Burak et al. 2004; Zalidis et al. 2002). Malta has a population of around 394,000, which triples each summer with the influx of tourists, who provide the main economic activity of the island. Average rainfall is 578 mm per annum. The country has no permanent rivers or lakes, and the water supply has traditionally been from the groundwater, which has become inadequate to support tourism. Malta has implemented a programme to desalinate sea water, and up to 70 per cent of Malta’s water now comes from desalination plants (Bremere et al. 2001).

Characteristics of the Present-day Climate In the western Mediterranean, the weather is principally dictated by proximity to the Atlantic Ocean and temperatures are less extreme than in the eastern Mediterranean. The highest daytime temperature recorded in Almeria, on the south-eastern coast of Spain, is 38◦ C, compared with 43◦ C in Athens. The lowest night time temperature recorded in Almeria is 0◦ C, compared with −4◦ C in Athens. Throughout the Mediterranean, winters are wet and summers are dry. Typically, between 70 and 80 per cent of rainfall is received between October and March, and around 40 per cent between December and February. Conditions become more extreme in the south-eastern Mediterranean and, in Alexandria, 98 per cent of rainfall is received between October and March, and 70 per cent between December and March (Figure 3.3b). Rainfall in the Mediterranean occurs in association with depressions, or centres of low pressure. As in middle and high latitudes, the development and steering of individual depressions is associated with the general thermal pattern and winds in the upper troposphere. The majority of depressions in the Mediterranean have their origins as lee depressions forming, for example, in the lee of the Alps and Atlas mountains (HMSO 1962). However, other mechanisms operate. In the Iberian Peninsula, it has been estimated that around half of rain-producing depressions are of Atlantic origin. These

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Fig. 3.2. Seasonal temperature and rainfall variations at selected sites around the Mediterranean.

Atlantic depressions weaken as they track across the peninsula and seldom bring rain to the Mediterranean basin. However, they may trigger the formation of new cyclones within the basin itself (Trigo et al. 1999, 2002). When waves of cold air, which accompany Atlantic depressions, encounter warm moist air over the

Mediterranean, the accompanying vertical instability, which may be enhanced in the presence of mountains, leads to the development of vigorous depressions, high rainfall, and high wind speeds (HMSO 1962). Thermal lows can also be important, for example, over the Iberian Peninsula (Chapter 18).

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(a)

1500 1000 750 500

Annual rainfall (mm/year)

250 0

0

500 km

0

500 km

(b)

7 5 3

Number of dry months

1 0 (in the north)

Fig. 3.3. (a) Mean annual rainfall and (b) length of the dry season across the Mediterranean basin (modified from Grenon and Batisse 1989).

The Climate System

The Influence of the Global Circulation Jet-streams in the Upper Atmosphere Jet streams are narrow, fast-flowing streams of air at about 11 km above the earth’s surface. Two westerly flowing jet streams can affect the Mediterranean region: the polar-front, most commonly found between latitudes 30 and 70◦ N, and the subtropical jet, between latitudes 20 and 50◦ N. Both are westerly flowing but follow a meandering course. Because they mark the boundary between cooler and warmer air masses, and because they can steer underlying surface depressions, their position at any point in time is important in determining the character of the surface weather. The polar-front jet stream generally lies north of the Mediterranean, but in winter it may be present and is then associated with outbreaks of cold air from the north (HMSO 1962). Abnormal summer heatwaves have been associated with a subtropical jet to the north of its normal position (Balafoutis and Makrogiannis 2001; Baldi et al. 2006).

Large-scale Pressure Patterns The Mediterranean is influenced by planetary-scale pressure patters that shift seasonally, moving equatorward in winter and poleward in summer as shown in Figure 3.1. This brings the Mediterranean basin under the influence of the subtropical high pressure belts in summer and under the influence of the westerly wind belts in winter. These large-scale synoptic controls are also discussed in relation to the nature of the marine environment in Chapter 2. This pattern dictates the large-scale average climate conditions which affect the Mediterranean. Superimposed on these planetary-scale average patterns, the behaviour of pressure centres over the Atlantic Ocean, and the large landmasses of Europe, Asia, and Africa influence the inter-annual and intra-seasonal variability of the regional climate. The behaviour of these pressure centres and their relationship with neighbouring centres (which may involve a see-sawing or oscillation of high pressure in one location and low pressure in another, known as a teleconnection pattern), is captured through the calculation of pressure indices. Some of the principal pressure indices and oscillations affecting the Mediterranean are described below.

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The North Atlantic Oscillation (NAO) has been found to influence rainfall amounts, especially in winter, in the north-western Mediterranean (Hurrell 1995) and, more remotely, in the eastern Mediterranean (Eshel and Farrell 2000; Eshel et al. 2000). The NAO is a measure of the track of storms and depressions across the North Atlantic Ocean and into Europe. A high NAO winter is one when the storm track is strong, with a north-eastward orientation taking depressions into north-west Europe. A low NAO winter has a weaker storm track with an east–west orientation taking depressions into Mediterranean Europe. The north-western Mediterranean will tend to have wetter winters when the NAO is weak and drier winters when the NAO is strong (Hurrell 1995). The eastern Mediterranean lies under continental influences, accounting for conditions more extreme than experienced in the west. The Indian Monsoon also plays a role in ensuring that the eastern Mediterranean is exceptionally dry in summer. The heating and associated uplift of air over India during the monsoon is linked to subsiding air and hence dry conditions in the eastern Mediterranean (Raicich et al. 2003). In the summer of 2002, Indian Monsoon rainfall was much below normal, and there was exceptionally heavy rain in many parts of southern and central Europe. It has been suggested that El Niño and La Niña events may affect Mediterranean climates. The Walker circulation is an east–west atmospheric circulation pattern of rising air above Indonesia and the western Pacific and sinking air above the eastern Pacific. Under normal conditions, Indonesia is wet and the eastern Pacific is dry. During El Niño events there is a weakening of the Walker circulation. During La Niña events the Walker circulation is especially strong. The effects of an El Niño period are directly linked to areas within the Pacific basin, causing droughts over Indonesia and intense rainfall events, flooding, and landslides in Peru. More importantly, its influence on seasonal weather conditions has been detected as far away as India, Africa, Antarctica, and North America (Bromwich et al. 2000). In the eastern Mediterranean there is some evidence that El Niño events are positively correlated with winter rainfall (Kadioglu et al. 1999; Price et al. 1998; Rodo et al. 1997). Other relationships between Mediterranean weather conditions and El Niño/La Niña events have been suggested, but these are generally weak or impermanent. Kutiel et al. (2002) found a teleconnection between pressure centres over the North Sea and the Caspian Sea. The North Sea Caspian Pattern Index (NCPI) is

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calculated from the North Sea–Caspian Sea pressure difference. Between October and April, when the NCPI is positive, temperatures are high over the Balkans, western Turkey, and the Middle East, with below average precipitation. A negative index during an extended winter period is associated with low temperatures and above-average precipitation in this region. Kutiel and Benaroch (2002) found that this index was a better predictor of the climate over the eastern Mediterranean region than the NAO.

The Mediterranean Oscillation So far we have considered the influence of pressure centres outside the Mediterranean basin. However, the elongated west–east shape of the Mediterranean Sea, and its size in relation to the global circulation, have led a number of authors to suggest that pressure differences between the western and eastern ends of the basin may influence climate variability at inter-annual scales. This teleconnection has been called the Mediterranean Oscillation and the Mediterranean Pressure Index (MPI). Here we call it the Mediterranean Oscillation (MO). It has variously been calculated from the pressure difference between Algiers and Cairo (e.g. Corte-Real et al. 1998), between Gibraltar and Israel (Lod Airport) (e.g. Palutikof 2003), and between Mersa Matruh (Egypt) and Marseilles (France) (e.g. Raicich et al. 2003). A number of different influences have been attributed to the MO/MPI. Maheras and Kutiel (1999) found that favourable circulation for high temperatures in the western basin (southerly flow) is associated with unfavourable circulation in the eastern basin (northerly flow) and vice versa. Piervitali et al. (1999) found high negative correlations between eastern Iberian and Italian rainfall and the MO/MPI, exceeding the relationship with the NAO throughout the year. Palutikof (2003) found significant negative relationships between the MO/MPI and rainfall over the western and central Mediterranean in winter and these were higher than the relationship between the NAO and rainfall in these locations. Strongly significant negative relationships between the MO/MPI and autumn rainfall were also found in the western basin. However, no relationship was found between the MO/MPI and rainfall in the eastern Mediterranean. Raicich et al. (2003) found a significant correlation between their MPI and the north–south wind component in the eastern Mediterranean (the Etesian wind regime), and hence negative correlations with Sahelian rainfall and the Indian monsoon.

Areas of Cyclogenesis In some preferred locations, interactions between the large-scale circulation and the complex geography create conditions ideal for the formation of cyclones. These locations, and the characteristics of the cyclones they produce, are described below and summarized in Table 3.1. The Mediterranean Sea is one of the most active regions of cyclogenesis in the world. Although systems are generally weak and shallow, occasionally strong, fast-moving, and active low pressure systems form, bringing very disturbed weather and severe conditions to the region. Strong cyclones may be associated with rainfall of up to 200 mm in 24 hours (Mahovic et al. 2005), and up to 800 mm in 24 hours has been recorded (Peñarrocha et al. 2002). Deep Mediterranean cyclones are commonly associated with high wind speeds, as shown in Chapter 18.

Cyclogenesis in the Northern Mediterranean Cyclones form in three principal areas of the northern Mediterranean: the Gulf of Genoa, the Aegean Sea, and the Black Sea (Figure 3.4). The cyclones produced are generally sub-synoptic in scale. They are triggered by the passage of remnant North Atlantic synoptic systems and their interactions with local topography. Cyclones may therefore form consecutively at the three centres as a single North Atlantic system passes each in turn (Trigo et al. 2002). Of these three centres, the best known is the Gulf of Genoa. Genoa cyclones, although most frequent in winter, may form throughout the year. They generally remain stationary to the south of the Alps. If they do move, they will follow one of two tracks. First, if there are anticyclonic conditions over the Balkans, Turkey, and the Black Sea, cyclones will move out of the Gulf of Genoa in a south-easterly direction towards the Ionian Sea. This provides ideal conditions for a Bora wind to develop as discussed below. Second, cyclones may move north-easterly across the Alps, and may bring extensive rainfall and catastrophic flooding to Austria, Germany, the Czech Republic, and Poland, as in the summer of 2002. Such a path may be associated with the development of Sirocco winds if the circulation of the low extends southwards into North Africa allowing air from the desert to move north. This second path is more commonly followed if cyclogenesis is

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TABLE 3.1. Centres of cyclogenesis in the Mediterranean Region The Adriatic and Ligurian Seas, including Gulf of Genoa Iberian Peninsula North Africa: Sahara Aegean Sea

Eastern Black Sea Cyprus

Middle East

Associated track

Associated mechanism

Seasonality

• SE-ward direction (Italy, Albania, Greece) • NE from the Adriatic into the Balkans Quasi-stationary

Lee-effect cyclogenesis and conditional instability, upper level vorticity

Declines in intensity towards summer

Creates intense rain across Buzzi and Tibaldi a large sector of the 1978 western basin. Associated with the Mistral

Thermally induced

Peaks June–August

Usually dry but may be thunderstorms Source of important spring rainfall, transported dust, and the Scirocco Increased storminess

• NE into the Med. Lee-effect cyclogenesis due • E along the African to Atlantic flow and coast towards Greece thermal instability NE towards the Black Conditional instability, Sea regenerated Genoan cyclones, lee-effect cyclogenesis NE into Europe Conditional instability, Aegean cyclogenesis E into the Middle East Lee-effect reintensification of western depressions

E and NE into Asia

Asian Monsoon mechanisms

Peaks May–June

Impacts

Declines in intensity towards summer strongest in January Declines in intensity Contributes to annual towards summer rainfall peak Increases in intensity Important source of rain towards summer (but also storms) for Cyprus, southern Turkey, and the Middle East Increases in intensity Dry and settled weather towards summer

References

Alonso et al. 1994 Egger et al. 1995

Flocas and Karacostas 1996 Radinovic 1987 Barry and Chorley 1992; Lagouvardos et al. 1996 Barry and Chorley 1992

Source: Harding (2006).

displaced eastwards from the Gulf of Genoa, into the Gulf of Venice. Aegean Sea cyclones are principally winter and spring phenomena. In the Black Sea, cyclones form throughout the year, reaching a maximum of one per week in July and August in the eastern Black Sea.

Cyclogenesis Over North Africa During late winter and early spring, North Africa becomes the primary region of cyclogenesis in the Mediterranean basin. The cyclones are usually dry, but are characterized by high winds close to their centres. They move extremely rapidly, following an eastward track south of the Atlas Mountains before moving over the Mediterranean Sea across the coast of Tunisia at or near the Gulf of Gabes. Their speeds out of North Africa may be as high as 20–25 ms−1 . Lee cyclogenesis is thought to play an important role in the formation of these cyclones. The synoptic situation favouring development is the presence of an upper trough lying over Spain with its axis lying NE–SW, producing a deep south-westerly flow over north-west Africa. North African lows are associated with strong easterly to south-easterly winds over the southern Mediterranean and high seas in the Strait of Sicily. On occasions they may be responsible for the entrainment and transfer of

dust into and across the Mediterranean from the Sahara Desert, as illustrated in Chapter 14.

Other Regions of Cyclogenesis These include: r The formation of thermal low pressure centres over

the Iberian Peninsula in summer (HMSO 1962).

r Winter cyclogenesis can be expected to take place

over Cyprus when a cold front approaches the Anatolian Plateau from the north. r In the Middle East, over Syria and Iraq, cyclogenesis may occur in summer as an extension of the Indian monsoon. Despite this cyclone activity, the upper troposphere of the eastern Mediterranean is dominated by strong descent as part of the Asian monsoon system, and conditions remain dry (Trigo et al. 1999).

Frequency of Cyclone Formation Cyclone counts will vary depending upon the definition used and the method of detection. However, it is widely accepted that the most prolific centre of cyclogenesis is the Gulf of Genoa. The UK Meteorological Office (HMSO

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Fig. 3.4. (a) Regions of cyclone genesis and dominant cyclone tracks in the Mediterranean (modified from HMSO 1962) and (b) a TERRA satellite image of a cyclone centred on the Ionian Sea, 24 February 2006. , accessed 20 November 2008).

The Climate System

1962) estimated the number of cyclones forming in an average year at this location at 60, compared to 51 in the Ionian Sea, 28 in the Middle East, and 14 in the Sahara. The principal areas of cyclogenesis and the tracks of depressions are shown in Figure 3.4a and a large winter cyclone centred on the Ionian Sea to the south of Italy is shown in Figure 3.4b.

Local Winds The complex geography of the Mediterranean basin creates many local wind systems, due to the interplay of the atmospheric circulation described above with the distribution of land and sea, and of mountains and coastal plains in close proximity. These local wind systems include the Levanter and Vendavales in the Strait of Gibraltar, the Bora in the former Yugoslavia and the northern Adriatic, the Sirocco over North Africa, and the Khamsin over Egypt into the eastern Mediterranean. Table 3.2 lists the local wind systems of the Mediterranean region. Many of these local winds are associated with cyclogenesis and the passage of depressions, as described above. Some can only form in the presence of low pressure systems whilst others require these systems to achieve their maximum development. Some local winds bring welcome relief in the summer heat. The wind known as the Etesian in Greece and the Metemi in Turkey forms when the presence of a high pressure system over Hungary and a low pressure system over Turkey channels a cool northerly airflow over Greece and Turkey in the summer. Others are regarded with fear. The Khamsin flows out of the desert bringing desiccating dust-laden air in late spring and early summer which can obscure visibility and irritate the eyes, nose, and mouth. Many local winds are katabatic, i.e. they are a downflow of higher-density cold air. The Bora forms when cold air ponds over the mountainous regions of the former Yugoslavia, eventually spilling over the high mountain passes and flowing down into the northern Adriatic. It is most common in winter and wind speeds can exceed 30 ms−1 . The Bora wind can also affect Venice in winter, causing damage and disrupting water traffic. The most severe local winds occur when the mesoscale atmospheric circulation interacts favourably with the regional topography, and the Bora is no exception. In Croatia, the Bora wind of 14/15 November 2004 was intensified by the presence of a deepening low pressure system over the southern Adriatic, and a strengthening of the central European anticyclone.

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It killed two people and injured over fifty. It disrupted energy supply and transport, and caused extensive damage to houses, harbours, and trees, with many olive trees uprooted. The Mistral is a katabatic cold wind of the western Mediterranean, most common in winter and spring. Cold air ponds over the Massif Central and Alps, and flows down the Garonne and Rhône valleys into the Gulf of Lion. Up to half of winter days in Marseilles can be affected by Mistral winds, which bring dry, cold, but sunny weather. The effects of the Mistral on the sea can sometimes be felt past Sicily, commonly causing wave heights of 5 to 6 m, and exceptionally up to 9 m. Although Mistral conditions can develop locally along the south coast of France, major Genoa cyclogenesis (see above) is necessary for an extensive Mistral to occur. Even after a Genoa low has moved eastwards, the Mistral can continue associated with a residual trough to the south of the Alps.

Extreme Events We have already described some severe wind events associated with local wind regimes in the Mediterranean. But, in people’s perception, the Mediterranean region is associated much more with other climaterelated hazards such as heatwaves and droughts, and with flash floods. Note that Mediterranean storms and floods are covered in greater detail in Chapter 18 in Part III of this volume on Natural Hazards.

Heatwaves Heatwaves in the Mediterranean have a number of impacts. They increase the number of deaths from heat stress, especially amongst the elderly, they increase energy demand for air conditioning, and they lead to widespread and devastating forest fires. Perhaps the most severe heatwave in the region was that of 2003, which affected the whole of Europe, and is considered to be the warmest summer since 1500 (Stott et al. 2004). The number of excess deaths in that heatwave around the Mediterranean is estimated at around 20,000 in Italy, 8,500 in Spain, and 2,000 in Greece. In Spain, mean daily temperatures of 33◦ C and over were recorded for at least half of the days (46/92 days) of the period June–August 2003 in 15 out of 48 cities. In 8 of these 15 cities, temperatures over 33◦ C were registered for more than 60 of the 92 days in the period (Simón et al. 2005). In Italy, daytime maximum temperatures remained between 38 and 40◦ C in most cities for weeks. As a result of the

78

Andrew Harding, Jean Palutikof, and Tom Holt TABLE 3.2. Local winds of the Mediterranean Location

Wind

France

Mistral

France

Bize

Spain

Levante

Spain

Leveche

Spain

Solano

Spain

Tramontana

Spain

Galerna

Spain

Cierzo

Italy

Maestro

Italy

Gregale

Balkans

Etesian

Balkans

Bora

Balkans

Varadarac

Africa

Khamsin

Africa

Scirocco

Africa

Ghibli

Characteristics • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • • •

Cold, dry, northerly or north-westerly katabatic flow Blows along the southern coast, as far as Genoa, most prevalent between Montpellier and Toulon Accompanied by clear weather and bright sunshine Associated with depressions in the Gulf of Genoa and a high pressure ridge from the Azores Cold, dry, northerly, north-easterly or north-westerly winter wind Blows in the mountainous regions of southern France in Languedoc Accompanied by heavy clouds Strong, moist, cool, north-easterly spring (Feb.–May) and autumn (Oct.–Dec.) wind Blows on the east and south coasts of Spain Associated with the Azores anticyclone Hot, dry, south-westerly wind Blows across south-east Spain Occurs in front of eastward moving depressions Hot, humid south-easterly or easterly summer wind Occurs on the east coast of Spain and in the Strait of Gibraltar Cold, dry, northerly; a form of Bora Blows along the northern Catalan coast and parts of Italy Cold, dry, north-westerly, all year but especially in winter Blows onto the north Spanish coast from the Atlantic Cold, dry, north-westerly, active for an extended (6-month) winter period Blows along the Ebro valley North-westerly summer wind Blows in the Adriatic when pressure is low over the Balkans Associated with fine weather and light clouds Strong, cold north-easterly wind (mainly winter) Blows in the Ionian Sea, usually lasting 2–3 days Frequently reaches gale force, usually accompanied by rain Northerly, summer, blows across the Aegean Associated with the Azores anticyclone and the Asian low Accompanied by high rates of evaporation and pleasant cool conditions Cold and dry, often gusty north-westerly winter katabatic wind Blows on the coast of Dalmatia into the Gulf of Trieste Associated with an eastward moving depression forced over the Balkans and cold air from Russia Similar to the Bora Blows along the Vardar valley into the Gulf of Thessaloniki Hot, dry, dusty, southerly Egyptian wind Associated with a trough moving across the eastern Mediterranean Accompanied by humidity drops to less than 10% Hot, very dry south to south-westerly wind that becomes moist as it traverses the Mediterranean Sea Comes from the desert climates of North Africa and the Near East, gathering moisture before reaching the Mediterranean • Blows along the majority of Mediterranean facing coasts, particularly of Algeria, Italy, and the Levant region • Very similar to the Scirocco but originating over Libya

Source: Harding (2006), modified from Rudloff (1981), Barry and Chorley (1992), and Kostopoulou (2003).

heatwave, and in the knowledge that such events are likely to become more common, many governments have put in place emergency plans to cope with future occurrences. It is likely that the 2003 heatwave is part of a rising trend, with a strong probability that this is related to global warming caused by human influence on the climate. Baldi et al. (2006) studied heatwaves between 1950 and 2000 in Italy; defining heatwaves as events that exceed the 90th percentile of the temperature

distribution (calculated over the period 1961–90) for six or more consecutive days. Table 3.3 shows the number of heatwaves per decade since 1950. It is clear that there has been a marked increase since 1990. Stott et al. (2004: 610) state that it is ‘very likely (confidence level >90%) that human influence has at least doubled the risk of a heatwave exceeding this threshold’. Half of all summers are expected to be as warm as 2003 by 2040, and the 2003 summer may approach the norm by 2080.

The Climate System TABLE 3.3. Decadal distribution of heatwave days Decade

Days (no.)

%

1951–60 1961–70 1971–80 1981–90 1991–2000

66 38 18 98 187

16 9 4 24 46

Total

407

100

Source: Modified from Baldi et al. 2006.

Droughts Although heatwaves and drought generally go hand in hand, the most severe droughts are measured in terms of their persistence. Thus in the summer of 2005 a severe drought affected the western Mediterranean region, but this was the culmination of dry conditions that began in the autumn of the previous year. In Spain, rainfall between November 2004 and March 2005, the traditional months of reservoir and groundwater recharge, was the lowest since 1947 (the first year of record). As reservoir stocks fell to 50 per cent of normal levels, rationing was imposed by regional governments in the east of Spain (García Herrera et al. 2007). Since rainfall is decreasing across much of the Mediterranean, it is likely that droughts are becoming more frequent. However, there is little evidence that this is the case. In part, this is because drought depends not only on the amount of rainfall, but also on the level of abstraction. Abstractions in the Mediterranean are increasing as water demand rises under pressure from competing uses for irrigation, tourism, industry, and domestic needs (Chapter 21).

Intense Rainfall Events Intense rainfall events are important because they cause flash floods and landslides, with damage to infrastructure and loss of life. In natural river catchments between about 25 to 2,500 km2 in area a flash flood could occur following a rainstorm of more than 200 mm in less than six hours. In urban areas flooding could occur in a built-up area of 1 to 100 km2 following even shorter rain storms of around 50 mm in less than one hour. Such intense rainfall amounts are generally produced by stationary meso-scale convective systems. Table 3.4 gives some examples of recent flash floods in the Mediterranean (Chapter 18). Over time, it appears that the contribution of intense rainfall events to the total precipitation input has increased in much of the Mediterranean. For example,

79

in Italy, daily rainfall amounts in excess of 128 mm contributed 4 per cent of total rainfall in the 1990s, compared to only 1 per cent in the 1950s. A similar increased contribution has occurred in Spain from rain events greater than 64 mm per day (Alpert et al. 2002).

Past Climates of the Mediterranean Long-term changes in climates of the Mediterranean are dealt with in several chapters elsewhere in this volume. In particular, Chapters 2 and 4 examine the long-term regional climate dynamics and external forcing factors during the Quaternary Period and earlier parts of the Cenozoic, and Chapter 9 looks in detail at Holocene climate and environmental change across the region. In order to set the context for the Mediterranean climates of the present and future, we look briefly at what has happened in the region over the last millennium and, in more detail, at trends in the last 50–100 years.

Mediterranean Climates of the Last Millennium The major climatic episodes of the last millennium in Europe are the Medieval Warm Period (MWP), the Little Ice Age (LIA), and Current Warm Period. Whereas the MWP and the LIA were due to natural variability, in the case of the LIA solar variability and volcanic activity, the Current Warm Period is widely considered to be due to human activities causing global warming (see below). The MWP lasted from the tenth to the fourteenth century. The LIA lasted from around the fourteenth to the end of the nineteenth century. Both the MWP and LIA are visible in the records of past climates from the Mediterranean. In north-west Spain, Martinez-Cortizas et al. (1999) found that the MWP was around 1.5◦ C warmer than the present day. In the eastern Mediterranean, it was a period of wetter conditions with, for example, high water levels in the Dead Sea and Sea of Galilee (Schilman et al. 2001). The LIA was a period of glacier advance in the Apennines and Pyrenees, and wintertime temperatures sometimes as much as 3◦ C colder than at present (Giraudi 2005). The impacts of the LIA on fluvial systems and glacial systems in the Mediterranean are presented in Chapters 11 and 12 respectively.

Climates of the Last Century We can trace the late-nineteenth-century warming that has taken place across the Mediterranean since the end

80

Andrew Harding, Jean Palutikof, and Tom Holt

of the LIA. Glaciers in the Pyrenees have shrunk substantially since that time. Total surface area has dropped from 40 km2 in 1895 to 30 km2 in 1958, 15 km2 in 1975, and just 8 km2 in 1992. In 1975 there were seventy glaciers in the Pyrenees but by 1992 this had shrunk to forty-one (Serrat and Ventura 1992). Today, the most southerly glacier in Europe is the Ghiacciaio del Calderone glacier, in the Apennines. Between 1884 and 1990, the area of this glacier shrank by half (D’Orefice et al. 2000) (Chapter 12). The Third Assessment of the Inter-governmental Panel on Climate Change stated that ‘the global average surface temperature has increased over the 20th century by about 0.6◦ C’ and that ‘there is new and stronger evidence that most of the warming observed over the last 50 years is attributable to human activities’ (IPCC Working Group I 2001). This warming is not uniform, nor is it constant over time. Using gridded temperature data, the IPCC found that warming has occurred over the Mediterranean in all seasons since the mid-1970s (Folland et al. 2001). However, more detailed analyses suggest more complex trends. Using station data for summer for the period 1950–99, Xoplaki (2002) found warming in the western and central Mediterranean, but cooling in the interior of eastern landmasses, especially eastern Turkey and the Balkans. Kutiel and Maheras (1998) studied data from 1873 to 1989 and found warming over the year as a whole, which was stronger in the western Mediterranean (around 0.4◦ C per century) and weaker in the eastern Mediterranean (only 0.2◦ C per century). When these data were analysed on a seasonal basis, a cooling trend was found in autumn in the eastern Mediterranean. Rainfall appears to have been largely decreasing across the Mediterranean over the last half century, although the patterns are complex. In Italy and Spain, this decrease reached 10–20 per cent over the 1951– 95 period (Piervitali et al. 1998; Romero et al. 1998). Winter rainfall amounts in the western and northern Mediterranean are related to the North Atlantic Oscillation (see above), with correlation coefficients of around –0.6 to –0.7 (Palutikof 2003). Between 1960 and 2000 the winter NAO followed a gradual rising trend from strongly negative values in the early 1960s to strongly positive values in the 1990s, and since then has generally remained positive. Since the Mediterranean receives most of its rainfall in the winter season, it is likely that the declining rainfall trend in the western and northern Mediterranean is related to the trend in the NAO. It is not possible to judge whether the unusual long-term trend in the NAO starting in the 1960s is in any way related to global warming. In contrast, some parts of the eastern

Mediterranean, particularly southern and central Israel, have shown an upward trend in rainfall over the last fifty years (Ben-Gai et al. 1994). Norrant and Douguédroit (2005) examined data from sixty-three rainfall stations across the Mediterranean to elucidate any trends between 1950 and 2000. They found a generally decreasing trend in rainfall, especially in the winter months, although not always statistically significant. Statistically significant declining trends in rainfall were found for: r r r r r r

March in the Atlantic region, October in Mediterranean Spain, December in the Gulf of Lion and Gulf of Genoa, January, winter, and the year in Greece, winter and the year in Italy, winter in the Near East.

A statistically significant increasing rainfall trend was found in April in the Gulfs of Genoa and Lion. Figure 3.5 shows composite temperature trends for the Mediterranean since 1960 for the year as a whole, and for winter and summer. These are calculated from eighty-four stations for the whole region, of which forty-one are in the western region, twenty-five in the centre and eighteen in the east (Harding 2006). These clearly show a warming trend in the west and central Mediterranean, which is present both seasonally and annually. This trend is absent in the eastern Mediterranean. Figure 3.6 shows composite rainfall trends for the same regions and time periods (ibid.). A declining trend is apparent in winter in all regions, whereas summer rainfall amounts are stationary. This confirms the findings of Norrant and Douguédroit (2005).

Future Climate Change Human influences are expected to lead to changes in climate and rising sea level during the twenty-first century and beyond. Industrial processes, the internal combustion engine, intensive agriculture, and deforestation all add greenhouse gases to the atmosphere, including carbon dioxide, the nitrous oxides, and methane. Whereas short-wave radiation from the sun is largely unaffected by these gases, long-wave radiation from the earth is absorbed. As atmospheric concentrations of these gases increase, so more long-wave terrestrial radiation is absorbed, leading to heating of the atmosphere and changes in pressure patterns and rainfall because the amount of energy in the atmosphere is increased.

Whole Mediterranean 4

Western Mediterranean 4

Annual

Central Mediterranean 4

Annual

Eastern Mediterranean 4

Annual

2

2

2

2

0

0

0

0

-2

-2

-2

-2

-4 1960 4

1970

1980

1990

-4 2000 1960 4

Winter (DJF)

1970

1980

1990

-4 2000 1960 4

Winter (DJF)

1970

1980

1990

-4 2000 1960 4

Winter (DJF)

2

2

2

2

0

0

0

0

-2

-2

-2

-2

-4 1960 4

1970

1980

1990

-4 2000 1960 4

Summer (JJA)

1970

1980

1990

-4 2000 1960 4

Summer (JJA)

1970

1980

1990

2000

4

Summer (JJA)

2

2

2

0

0

0

0

-2

-2

-2

-2

1980 Year

1990

-4 2000 1960

1970

1980 Year

1990

-4 2000 1960

1970

1980 Year

1990

2000

1980

1990

2000

1970

1980

1990

2000

1990

2000

Summer (JJA)

-4 1960

1970

1980 Year

Fig. 3.5. Composite graphs of annual, winter (DJF), and summer (JJA) temperature in the whole Mediterranean and the western, central, and eastern basins, 1960–2000. Units are standard deviations.

The Climate System

1970

1970

Winter (DJF)

-4 1960

2

-4 1960

Annual

81

82

Western Mediterranean 4

Annual

Central Mediterranean 4

Annual

Eastern Mediterranean 4

Annual

2

2

2

2

0

0

0

0

-2

-2

-2

-2

-4 1960 4

1970

1980

1990

2000

-4 1960 4

Winter (DJF)

1970

1980

1990

2000

-4 1960 4

Winter (DJF)

1970

1980

1990

-4 2000 1960 4

Winter (DJF)

2

2

2

2

0

0

0

0

-2

-2

-2

-2

-4 1960 4

1970

1980

1990

2000

-4 1960

1970

1980

1990

2000

1970

1980

1990

2000

1970

-4 1960

1970

Summer (JJA)

Summer (JJA)

Summer (JJA)

2

2

2

0

0

0

0

-2

-2

-2

-2

1970

1980 Year

1990

2000

-4 1960

1970

1980 Year

1990

2000

-4 1960

1990

2000

1980

1990

2000

1980 Year

1990

2000

Winter (DJF))

2

-4 1960

1980

4

4

4

Summer (JJA)

-4 1960

Annual

1970

1980 Year

1990

2000

-4 1960

1970

Fig. 3.6. Composite graphs of annual, winter (DJF), and summer (JJA) precipitation in the whole Mediterranean and the western, central, and eastern basins, 1960–2000. Units are standard deviations.

Andrew Harding, Jean Palutikof, and Tom Holt

Whole Mediterranean 4

The Climate System

The present-day climate of the Mediterranean already challenges economic activities in the region and puts ecosystems at risk. For example: r Rainfall barely provides sufficient water to support

many ecosystems and economic activities. Expensive irrigation systems and water transfer systems are required to maintain agricultural productivity in countries such as Spain. r Temperatures are high in summer, and any further warming could lead to heat stress, the need for air conditioning, and outdoor conditions too uncomfortable to sustain beach tourism. r Increases in damaging extremes (e.g. flash floods, droughts, and wildfires) could lead to further loss of life and economic damages. r Some coastal areas of the Mediterranean are already affected by flooding from the sea, particularly along the eastern coast of Italy. Venice experiences flooding due to a combination of subsidence and a rise in the levels of the lagoon waters. It is expected that the city will be uninhabitable at the end of this century unless new methods of protection from the water are installed. It is therefore very important to understand how future climates might change in the Mediterranean in response to global warming, and what the local impacts could be (Jeftic et al. 1992, 1996). The Inter-governmental Panel on Climate Change (IPCC) Third Assessment Report looked at future climate change at the regional scale. For the Mediterranean it found that: r The Mediterranean region is expected to warm at a

rate greater than global mean warming. Warming in the summer season (June to August) is expected to be as much as 40 per cent above the global mean. Taking into account the range of uncertainty, this implies a summer warming of at least 2◦ C, and possibly as much as 6◦ C by 2100. r Little change in winter-time rainfall is expected although summers are expected to become considerably drier (Giorgi et al. 2001). These two facts are not unrelated. Summer drying will mean lower evaporation rates. Evaporation requires latent heat; therefore energy which would be required to enable evaporation to take place will be available for warming the air. Summer warming is therefore predicted to be considerably above the global mean. Globally, sea level is likely to increase by anything between 30 cm and 50 cm by 2100, although the rise could be anything between 9 cm and 88 cm because of

83

the uncertainties involved in making the estimate (IPCC Working Group I 2001). The regional changes will be different, depending on local land movements (Chapter 13). However, sea-level rise is a threat to coastal regions of the Mediterranean including Venice. All areas of high population density and/or heavy infrastructure development are vulnerable, especially where the land is subsiding. Such a location is the Nile Delta, where many millions of people are involved in farming, and fishing in the coastal lagoons. Their livelihoods would be severely threatened by sea-level rise.

Local Changes in Mediterranean Climates So far we have looked only at Mediterranean-wide changes. However, we can examine regional variations using climate models; these are computer-based simulations of the behaviour of the atmosphere, oceans, land surface, and ice, and their interactions. By changing the concentrations of greenhouse gases in these models, future changes in climate can be simulated. To explore future climate changes in the Mediterranean we use output from the Hadley Centre Regional Climate Model, HadRM3 (Hudson and Jones 2002). It is important to note that these simulations are just one realization of a wide range of possible futures. Such a realization is generally called a scenario, or plausible future climate. Different assumptions about future economic activities, and hence future emissions of greenhouse gases and their concentrations in the atmosphere, can be made. The Special Report on Emissions Scenarios (Nakicenovic and Swart 2000) defined four future pathways of economic and demographic development up to 2100, with associated emissions scenarios. Here, we look at the A2 pathway, which defines a future world of very rapid economic growth, low population growth, and rapid introduction of new and more efficient technology. Major underlying themes are economic and cultural convergence and capacity-building, with a substantial reduction in regional differences in per capita income. Figure 3.7 shows plausible future changes in winter and summer mean temperature and rainfall across the Mediterranean in 2070–99 compared to 1961–90, based on the HadRM3 model using an A2 emissions scenario. Winter temperature increases of up to 5◦ C are simulated over land, with the largest increases over central Europe, Turkey, and the Middle East, and the lowest changes (2–3◦ C) over the maritime coastline of Europe. Summer temperature increases are larger, up to 7◦ C by

84

Andrew Harding, Jean Palutikof, and Tom Holt

50°N

5

50°N

60 40

4

20 3

40°N

40°N 0

2 1

30°N 10°W

0°

10°E

20°E

30°E

-20

10°W

40°E

50°N

-40

30°N

7

0°

10°E

20°E

30°E

40°E

50°N

0

6

-10

5 4

40°N

10°W

0°

10°E

20°E

30°E

-30

3

-40

2

-50

1

30°N

-20 40°N

40°E

-60

30°N 10°W

0°

10°E

20°E

30°E

40°E

Fig. 3.7. Future changes in temperature (left-hand side) and precipitation (right-hand side) over the Mediterranean (2070–99 minus 1961–90). Winter (DJF) changes are shown above, and summer (JJA) changes are shown below. Units are ◦ C for temperature and mm for precipitation. Changes are shown for the SRES A2 scenario in the Hadley Centre Regional Climate Model HadRM3—see text for further explanation.

2070–99, and greatest over the northern shore of the Mediterranean. Spatial variations of precipitation changes are, as expected, greater and less coherent than for temperature. In winter, the largest reductions in rainfall are seen over the Mediterranean coast and islands, extending inland to include much of Greece, Turkey, and southern and central Spain. Modelled reductions in average winter rainfall are generally of the order of −20 mm over the three month period in these areas, but can be up to −40 mm in isolated areas in southern Turkey and Greece. Increased rainfall is suggested over the northern shore of the Adriatic, into the Alps and central and northern France, with increases as great as +60 mm over the Alps. In summer, rainfall is expected to remain unchanged or decrease, with no areas of increased rainfall indicated. The greatest reductions are seen over the Alps, in a belt which extends southwest to include the Pyrenees, to as much as −60 mm over three summer months. Smaller declines are projected over North Africa and the Middle East.

Future Changes in Extremes In a region that can experience several rain-free months per year, and where intense rainfall events can lead to

devastating flash floods, any future changes in these extremes could be very damaging. Climate models such as HadRM3 can provide the necessary information to explore future changes in extreme events. Figure 3.8(a) shows the changes suggested by HadRM3 in the length of the summer drought in the Mediterranean. Most places around the Mediterranean experience a prolonged period of summer drought every year, with rainfall being an exceptional occurrence in August. Any tendency for this dry period to prolong will have profound implications for water resources in the region. Figure 3.8(a) suggests that, in future, the summer drought period will lengthen in most land areas of the Mediterranean south of 40◦ N, typically by around ten days. This tendency is shown to be particularly severe over the Middle East and Egypt, reaching an increase of twenty to thirty days. North of 40◦ N, most areas show no change or a decrease in the length of the summer drought, by as much as ten to twelve days. Figure 3.8(b) shows the future change projected by HadRM3 in the maximum amount of rainfall received in a single five-day period. The spatial patterns are extremely fragmented, but throughout Spain, the southern shore of the Mediterranean, and the Middle East, the maximum five-day rainfall is suggested to remain unchanged or decrease. The principal land areas

The Climate System

85

(a) 50°N

30 20 10

40°N

0 -10

30°N 10°W

(b)

0°

10°E

20°E

30°E

40°E

50°N 40

20

0

40°N

-20

-40

30°N 10°W

0°

10°E

20°E

30°E

40°E

Fig. 3.8. Future changes in (a) length of the summer drought, in days; (b) maximum five-day precipitation, in mm, over the Mediterranean (2070–99 minus 1961–90). Changes are shown for the SRES A2 scenario in the Hadley Centre Regional Climate Model HadRM3—see text for further explanation.

where large increases are suggested are central Italy, the former Yugoslavia, and Switzerland, where increases of 30–40 mm are simulated.

Conclusions We have looked in this chapter at all aspects of climate over the Mediterranean in the recent past, the present, and future and we have outlined the major controlling factors in relation to the global climate system. In one chapter it is not possible to treat every aspect of climate in great detail. Those who seek a more detailed

treatment are directed to Mediterranean Climate Variability (Lionelli et al. 2006) and to a special issue of Global and Planetary Change entitled Mediterranean: Trends, Variability and Change (Piero et al. 2008). Here, the goal has been to present the importance of the climate of the Mediterranean to all aspects of life in the region. This importance derives from, first, rainfall amounts which are barely sufficient and, second, the risks associated with weather extremes. Rainfall amounts, and the strong seasonality of rainfall, mean that competition for water is intense, and a limiting factor for regional development. In Catalonia,

86

Andrew Harding, Jean Palutikof, and Tom Holt TABLE 3.4. Summary of some recent flash floods in Mediterranean Europe

Place

Date

Barcelona area, Spain Nîmes, France

25/9/1962

Vaison-la-Romaine, France Brig, Switzerland

26/9/1992

Versilia, Italy

19/6/1996

Biescas, Spain

7/8/1996

Corbières, France

12–13/11/1999

Sovearto, Italy

9–10/9/2000

Anduze, France

8–9/9/2002

4–5/10/1988

22–4/9/1993

Event

Loss of life

250 mm of rainfall in 2 hours caused flash flooding in the Besos River basin Urban flood from the Cadereaux watersheds (less than 50 km2 ). Peak flow in the city of about 1,000 m3 /s Over 300 mm of rainfall in 24 hours; peak flow of about 1,000 m3 /s at Vaison-la-Romaine 40 mm of rainfall in 24 hours on 23rd and 65 mm on 24th and maxima at Simplon (in Saltina River catchment above Brig) of 120 mm on 23rd and 220 mm on 24th. Flooding in Brig up to 3 m for 12 hours 400 mm of rain in less than 6 hours with maximum rainfall intensity of about 88 mm in 30 mins Over 250 mm of rainfall in 6 hours. Peak flow of 400–600 m3 in small upstream catchment of the Aras River (area 18 km2 ) devastated camping site located at outlet Rural flood of the Aude River (4,840 square km), 30–50% of peak flow being produced by a 123-km2 watershed (Guame et al. 2004) 350 mm of rain in 24 hours with a peak at 185 mm in 6 hours, causing landslides Over 600 mm of rain in 24 hours at Anduze; peak flow of the Gard River of 5,000 to 7,000 m3 /s for a 1,400 km2 basin

More than 1,000 deaths 11 deaths

Insured costs US $80 million €610 million

58 deaths 2 deaths

US $33 million

26 deaths

US $33 million

87 deaths

35 deaths

US $3 million

16 deaths (12 from a campsite) 25 deaths (in the same river 35 people were killed in 1958)

€1.2 billion

Source: http://natural-hazards.jrc.ec.europa.eu/activities_flood_flashflood.html accessed 31st March 2009.

Spain, 4.5 million people are affected by chronic water shortages, and the authorities are pressing for the construction of a pipeline to divert water from the Rhône in France to Barcelona. At the other end of the Mediterranean, Turkey has been accused by Syria and Iraq of depriving them of much-needed water, as it continues to build a series of dams along the Euphrates and Tigris. Turkey also plans to sell water from the Manavgat River (which flows south from the Anatolian Plateau into the Mediterranean) across the Middle East (Gruen 2000). The regional climate is characterized by extremes, which cause economic damage costing many millions of Euros and, in the most severe cases, fatalities. Large inter-annual variability is linked to the occurrence of extreme events such as droughts. Intense rainfall events, lasting only a few hours, are also a feature of the region’s climates (Table 3.4). Against this background of marginal water resources and damaging extremes, any tendency for the Mediterranean climate to change in future, whether for the better or the worse, is of intense interest.

Acknowledgements Original research presented in this chapter was carried out as part of the project ‘Modelling the Impact of Climate Extremes’

funded by the European Union under contract EVK20CT20010018 and a University of East Anglia research studentship. Climate model data was supplied by the Climate Impacts LINK Project (DEFRA Contract EPG 1/1/154) on behalf of the Hadley Centre and UK Met Office. We also thank Jamie Woodward and the external reviewer for reviewing the original manuscript.

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The Climate System Barry, R. G. and Chorley, R. J. (1992), Atmosphere, Weather and Climate. Routledge, London. BBC (British Broadcasting Corporation) (n.d.). World Water Crisis. , accessed 17 April 2006. (n.d.). BBC Weather Country Guide to Egypt, , accessed 7 September 2007. Ben-Gai, T., Bitan, A., Manes, A., and Alpert, P. (1994), Longterm changes in annual rainfall patterns in southern Israel. Theoretical and Applied Climatology 49: 59–67. Bremere, I., Kennedy, M., Stikker, A., and Schippers, J. (2001), How water scarcity will affect the growth in the desalination market in the coming 25 years. Desalination 138: 7–5. Bromwich, D. H., Rogers, A. N., Kallberg, P., Cullather, R. I., White, J. W. C., and Kreutz, K. J. (2000), ECMWF analyses and reanalyses depiction of ENSO signal in Antarctic precipitation. Journal of Climate 13: 1406–20. Burak, S., Dogan, E., and Gazioglu, C. (2004), Impact of urbanization and tourism on coastal environment. Ocean and Coastal Management 47: 515–27. Buzzi, A. and Tibaldi, S. (1978), Cyclogenesis in the lee of the Alps: A case study. Quarterly Journal of the Royal Meteorological Society 104: 271–87. Corte-Real, J., Sorani, R., and Conte, M. (1998), Climate change, in P. Mairota, J. B. Thornes, and N. Geeson (eds.), Atlas of Mediterranean Environments in Europe: The Desertification context. John Wiley, Chichester, 34–6. D’Orefice, M., Pecci, M., Smiraglia, C., and Ventura, R. (2000), Retreat of Mediterranean glaciers since the Little Ice Age: Case study of Ghiacciaio del Calderone, central Apennines, Italy. Arctic, Antarctic, and Alpine Research 32: 197–201. Egger, J., Alpert, P., Tafferner, A., and Ziv, B. (1995), Numerical experiments on the genesis of Sharav cyclones: idealised simulations. Tellus Series A: Dynamic Meteorology and Oceanography 47: 162–74. Eshel, G. and Farrell, B. F. (2000), Mechanisms of Eastern Mediterranean rainfall variability. Journal of the Atmospheric Sciences 57: 3219–32. Cane, M. A., and Farrell, B. F. (2000), Forecasting Eastern Mediterranean drought. Monthly Weather Review 128: 3618–30. Flocas, H. J. and Karacostas, T. S. (1996), Cyclogenesis over the Aegean Sea: identification and synoptic categories. Meteorological Applications 3: 53–61. Folland, C. K., Karl, T. R., et al. (2001), Observed climate variability and change, in J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden, and D. Xiaosu (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group 1 to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge. García-Herrera, R., Paredes, D., Trigo, R. M., Trigo, I. F., Hernández, E., Barriopedro, D., and Mendes, M. A. (2007), The outstanding 2004/05 drought in the Iberian Peninsula: associated atmospheric circulation. Journal of Hydrometeorology 8: 483–98. Gaume, E., Livet, M., Desbordes, M., and Villeneuve, J. P. (2004), Hydrological analysis of the River Aude, France, flash flood on 12 and 13 November, 1999. Journal of Hydrology 286: 135–54. Giorgi, F., Hewitson, B., et al. (2001), Regional climate information—evaluation and projections, in J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden, and

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D. Xiaosu (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group 1 to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge. Giraudi, C. (2005), Middle to Late Holocene glacial variations, periglacial processes and alluvial sedimentation on the higher Apennine massifs (Italy). Quaternary Research 64: 176–84. Grenon, M. and Batisse, M. (1989), Futures for the Mediterranean Basin: The Blue Plan. Oxford University Press, Oxford. Gruen, G. E. (2000), Turkish waters: source of regional conflict or catalyst for peace? Water, Air and Soil Pollution 123: 565–79. Harding, A. E. (2006), Changes in Mediterranean climate extremes: Patterns, causes and impacts of change. Doctoral Dissertation Ph.D., Climatic Research Unit, University of East Anglia. HMSO, Meteorological Office (1962), Weather in the Mediterranean, i. General Meteorology. HMSO, London. Hudson, D. A. and Jones, R. G. (2002), Regional Climate Model Simulations of Present-Day and Future Climates of Southern Africa, Hadley Centre Technical Note 39. Hadley Centre, Met Office, Exeter. Hurrell, J. W. (1995), Decadal Trends in the North-Atlantic Oscillation—Regional Temperatures and Precipitation. Science 269: 676–9. IPCC Working Group I (2001), Summary for Policymakers, in J. T. Houghton, Y. Ding, D. J. Griggs, M. Noguer, P. J. van der Linden, and D. Xiaosu (eds.), Climate Change 2001: The Scientific Basis. Contribution of Working Group 1 to the Third Assessment Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge. Jeftic, L., Milliman, J. D., and Sestini, G. (1992), Climate Change in the Mediterranean, Edward Arnold, London. Keskes, S., and Pernetta, J. C. (1996), Climatic Change and the Mediterranean, ii. Edward Arnold, London. Kadioglu, M., Tulunay, Y., and Borhan, Y. (1999), Variability of Turkish precipitation compared to El Nino events. Geophysical Research Letters 26/11: 1597–600. Köppen, W. (1936), Das geographische System der Klimate, in W. Köppen and R. Geiger (eds.), Handbuch der Klimatologie, iii Gebrüder Borntraeger, Berlin. Kostopoulou, E. (2003), The relationships between atmospheric circulation patterns and surface climatic elements in the eastern Mediterranean. Doctoral Dissertation Ph.D., Climatic Research Unit, University of East Anglia. Kutiel, H. and Benaroch, Y. (2002), North Sea–Caspian Pattern (NCP)—an upper level atmospheric teleconnection affecting the Eastern Mediterranean: Identification and definition. Theoretical and Applied Climatology 71: 17–28. and Maheras, P. (1998), Variations in the temperature regime across the Mediterranean during the last century and their relationship with circulation indices. Theoretical and Applied Climatology 61: 39–53. Turkes, M., and Paz, S. (2002), North Sea Caspian Pattern (NCP)—an upper level atmospheric teleconnection affecting the eastern Mediterranean—implications on the regional climate. Theoretical and Applied Climatology 72: 173–92. Lagouvardos, K., Kotroni, V., Dobricic, S., Nickovic, S., and Kallos, G. (1996), The storm of October 21–22, 1994, over Greece: Observations and model results. Journal of Geophysical Research-Atmospheres 101(D21), 26217–26.

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Lionelli, P., Malanotte-Rizzoli, P., and Boscolo, R. (2006), Mediterranean Climate Variability, Developments in Earth and Environmental Sciences 4. Elsevier, Amsterdam. Maheras, P. and Kutiel, H. (1999), Spatial and temporal variations in the temperature regime in the Mediterranean and their relationship with circulation during the last century. International Journal of Climatology 19: 745–64. Mahovic, N. S., Drvar, D., and Vrsnak, D. P. (2005), EUMeTrain: Cyclogenesis in the Mediterranean, , accessed 7 September 2006. Martinez-Cortizas, A., Pontevedra-Pombal, X., Garcia-Rodeja, E., Novoa-Muñoz, J. C., and Shotyk, W. (1999), Mercury in a Spanish peat bog: Archive of climate change and atmospheric metal deposition. Science 284: 939–42. Nakicenovic, N. and Swart, R. (2000), Emissions Scenarios. Special Report of the Intergovernmental Panel on Climate Change. Cambridge University Press, Cambridge. Norrant, C. and Douguédroit, A. (2005), Monthly and daily precipitation trends in the Mediterranean (1950–2000). Theoretical and Applied Climatology 83: doi 10.1007/s00704-0050163-y. Palutikof, J. (2003), Analysis of Mediterranean Climate Data: Measured and Modelled, in H.-J. Bolle, Mediterranean Climate: Variability and Trends, Springer, Berlin, 125–32. Peñarrocha, D., Estrela, M. J., and Millán, M. (2002), Classification of daily rainfall patterns in a Mediterranean area with extreme intensity levels: the Valencia region. International Journal of Climatology 22: 677–95. Piero, L., Serge, P., and Xavier, R. (eds.) (2008), Mediterranean climate: trends, variability, and change. Global and Planetary Change 63/2–3: 87–282. Piervitali, E., Colacino, M., and Conte, M. (1998), Rainfall over the Central-Western Mediterranean basin in the period 1951–1995. Part I: Precipitation trends. Nuovo Cimento Della Societa Italiana Di Fisica C-Geophysics and Space Physics 21: 331–44. Conte, M., and Colacino, M. (1999), Rainfall over the Central-Western Mediterranean basin in the period 1951– 1995. Part II: Precipitation scenarios. Nuovo Cimento Della Societa Italiana Di Fisica C-Geophysics and Space Physics 22: 649–61. Price, C., Stone, L., Huppert, A., Rajagopalan, B., and Alpert, P. (1998), A possible link between El Nino and precipitation in Israel. Geophysical Research Letters 25: 3963–6.

Radinovic, D. (1987), Mediterranean cyclones and their influence on the weather and climate. PSMP Report Series 24. World Meteorological Organisation, Geneva. Raicich, F., Pinardi, N., and Navarra, A. (2003), Teleconnections between Indian monsoon and Sahel rainfall and the Mediterranean. International Journal of Climatology 23: 173–86. Rodo, X., Baert, E., and Comin, F. A. (1997), Variations in seasonal rainfall in southern Europe during the present century: Relationships with the North Atlantic Oscillation and the El Nino Southern Oscillation. Climate Dynamics 13: 275–84. Romero, R., Guijarro, J. A., Ramis, C., and Alonso, S. (1998), A 30-year (1964–1993) daily rainfall data base for the Spanish Mediterranean regions: First exploratory study. International Journal of Climatology 18: 541–60. Rudloff, W. (1981), World Climates. Stuttgart, Wissenschaftliche Verlagsgesellschaft. Schilman, B., Bar-Matthews, M., Almogi-Labin, A., and Luz, B. (2001), Global climate instability reflected by Eastern Mediterranean marine records during the late Holocene. Palaeogeography, Palaeoclimatology, Palaeoecology 176: 157–76. Serrat, D. and Ventura, J. (1992), Glaciers of the Pyrenees, Spain and France. US Geological Survey Professional Paper 1386-E (Glaciers of Europe). Simón, F., Lopez-Abente, G., Ballester, E., and Marti, F. (2005), Mortality in Spain during the heatwaves of summer 2003. Euro Surveillance 10: 156–61. Stott, P. A., Stone, D. A., and Allen, M. R. (2004), Human contribution to the European heatwave of 2003. Nature 432: 610–14. Trigo, I. F., Bigg, G. R., and Davies, T. D. (2002), A climatology of cyclogenesis mechanisms in the Mediterranean. Monthly Weather Review 130: 549–69. Davies, T. D., and Bigg, G. R. (1999), Objective climatology of cyclones in the Mediterranean Region. Journal of Climate 12: 1685–96. Xoplaki, E. (2002), Climate variability over the Mediterranean. Doctoral Dissertation Ph.D., Institute of Geography, University of Bern. Zalidis, G., Stamatiadis, S., Takavakoglou, V., Eskridge, K., and Misopolinos, N. (2002), Impacts of agricultural practices on soil and water quality in the Mediterranean region and proposed assessment methodology. Agricultural Ecosystems and Environment 88: 137–46.

This chapter should be cited as follows Harding, A. E., Palutikof, J., and Holt, T. (2009), The climate system, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 69–88.

4

Cenozoic Climate and Vegetation Change Chronis Tzedakis

Introduction This chapter traces the history of vegetation change in the Mediterranean area in response to climate variability over the last 65 million years (Myr 1 ), with particular emphasis on the most recent part of the record. Compared to other continental areas of the globe, the Mediterranean region is somewhat unusual in the abundance of palaeobotanical information (especially palynological) that is available. This is a function of its geological setting, which in some cases has led to the relatively undisturbed accumulation of thick sedimentary sequences in tectonic and volcanic basins. These sequences have provided an opportunity to develop long records of vegetation change, sometimes extending over hundreds of thousands of years. In the marine realm, sedimentary records from the Mediterranean Sea are not only providing palaeoceanographic information but also beginning to yield palynological information, which can be placed directly within a chronological and palaeoclimatic framework. However, it is only in the uppermost part of the geological column (i.e. in the Quaternary) that there are enough records to construct a continuous thread of vegetation changes and allow meaningful comparisons between sites to determine local differences and transregional similarities (e.g. Magri et al. 2004). Moreover, the majority of terrestrial records extending before the Holocene are located in southern Europe, 1

‘Ma’ and ‘ka’ are used here to denote million and thousand years before present, respectively, while ‘Myr’ and ‘kyr’ are reserved for durations. Unless otherwise stated, calendar years are used.

while the coverage of the Near East is low and of North Africa even lower (Figure 4.1). The information available for earlier periods anywhere in the Mediterranean is fragmentary at best, with large parts of the record not represented. This means that despite the relative wealth of information, the palaeobotanical record from the Mediterranean region remains very much incomplete, with significant temporal and geographical gaps. Thus, instead of providing a linear narrative of the last c.65 Myr, the approach followed here is to structure this review into separate sections, each representing different scales of environmental variability, and attempt to establish the general pattern of vegetation responses to it. These environmental regimes are here defined as (1) mega-scale (long-term climate trends), (2) macroscale (orbitally driven (Milankovitch) changes), (3) meso-scale (sub-orbital variability), and (4) micro-scale (interannual variability) (Table 4.1). The first category refers to a shifting climatic mean forced by global plate tectonics, leading to a gradual transformation of continental and oceanic palaeogeography and changes in atmospheric greenhouse gases (Zachos et al. 2001). Oscillating about this mean are higherfrequency climate changes generated by variations in the Earth’s orbital geometry that affect the seasonal and latitudinal distribution of incoming solar radiation (Milankovitch 1930; Hays et al. 1976). Superimposed on these trends and rhythms are abrupt climate changes lasting centuries to millennia, whose origin is still debated but whose effects are becoming increasingly better documented (e.g. Voelker 2002; McManus et al. 1999). Finally, the last category refers to annual

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Ljubljan a

Bouchet/Ribain s

Lagaccion e

La B ord e

V. di Castiglione Banyole s

T. Philippo n

Monticchi o

40°N

Ioannina

Sile s MD95-2042

Pa dul

ODP976

Xinias

Va n

Ko pais

San Rafae l Dar Fatma

MD95-2043

Ghab

MD84-627

Tigalmin e Tigalmamine

Ma’ale Efrayim

Sore q

30° 0°

20°

40°E

Fig. 4.1. Location of sites discussed in the text.

TABLE 4.1. Scales of environmental variability and vegetation responses Environmental regimes Mega-scale (106–7 yr)

Macro-scale (104–5 yr) Meso-scale (102–3 yr) Micro-scale (100–1 yr)

Vegetation responses

Tectonic transformation of continental and oceanic palaeogeography, long-term changes in atmospheric and oceanic circulation patterns and changes in atmospheric concentration of greenhouse gases. Long-periodic orbital variations Orbitally driven (Milankovitch) climate variability, ice sheet build-up and decay Millennial/centennial sub-orbital climate variability Interannual climate variability, fire, volcanic eruptions, pathogenic attacks

to decadal changes that, however, are not considered here because the vast majority of the palaeorecords do not have the necessary stratigraphic detail to resolve them. Moreover, this chapter does not discuss the (mainly Holocene) record of anthropogenic impact on vegetation (see Part II), and concentrates instead on examining the vegetation response to natural climate variability. In the following sections, the global and Mediterranean environmental backgrounds and the response of vegetation are examined under each of the first three scales. In the final section, an attempt is made to consider the effects of interactions between the scales, on the duration of forested periods, the amplitude of suborbital-scale changes, genetic divergence, and the origin

Evolutionary changes (adaptation, speciation, extinction). Appearance of new biomes

Glacial–interglacial vegetation cycles (ecosystem change, individualistic response, refugia, migration, extirpations, divergence) Population contraction/expansion, replacement Disturbance, population collapse

of the mediterranean 2 climate rhythm and broadleaved evergreen sclerophylls.

Mega-scale (106–7 yr) Global Changes: Climate Plate tectonics is the main driver of the climate shifts observed at this scale leading to the gradual and mostly 2 The use of the term ‘Mediterranean’ (with initial capital) is reserved for the actual geographical region. The word begins in lower case when denoting mediterranean-type climates or vegetation that are not necessarily exclusive to this geographical area.

Cenozoic Climate and Vegetation Change

unidirectional modification of the Earth’s boundary conditions (Zachos et al. 2001). Most prominent amongst the changes over the last 65 Myr is the widening of the North Atlantic, the opening of the Antarctic gateways (Tasmanian and Drake passages), the collision of the Indian and Asian plates and uplift of the Himalayas and Tibetan Plateau, the emergence of the Isthmus of Panama and the decline in atmospheric CO2 (see review by Zachos et al. 2001 and references therein; Pagani et al. 2005). These changes contributed to long-term global cooling and cryospheric development, which, however, did not proceed gradually and uniformly but was punctuated by a series of steps, representing rapid transitions into new climate states (Kennett 1995). In addition, long-period orbital variations, such as the 1.2 Myr cycle in obliquity and the 2.3 Myr cycle in eccentricity, modulating the amplitude of higherfrequency astronomical changes, could also be included in this category (Lourens and Hilgen 1997). Indeed, some of the transitions into new climate states are associated with changes in the long-period orbital variations, suggesting that the Earth System may be more susceptible to change during such nodal points (e.g. Lourens and Hilgen 1997; Wade and Pälike 2004) Figure 4.2 shows a compilation of benthic oxygen isotope (‰18 Obenthic ) records by Zachos et al. (2001), which provides a global view of Cenozoic climate change. The first part of the Cenozoic was characterized by ice-free conditions. Warmest conditions were reached during the Early Eocene, followed by a cooling trend during which deep-sea temperatures dropped from 12 to 4.5◦ C. A prominent and rapid shift in ‰18 O values at 34 Ma signals the onset of major ice accumulation in East Antarctica (Shackleton and Kennett 1975). Ice sheets estimated at c.50 per cent of that of the present-day ice sheet persisted in Antarctica until 26 Ma. Atmospheric CO2 reconstructions (Pagani et al. 2005) show values fluctuating between 1,000 and 1,500 ppmv during the Middle to Late Eocene, with an overall downward trend. By 34 Ma, CO2 concentrations had reached levels sufficiently low to have triggered the initiation of ice accumulation in Antarctica, while by the latest Oligocene, CO2 concentrations had reached pre-industrial levels of ∼290 ppmv. In general, the co-evolution of ‰18 O records and CO2 concentrations during the Eocene and Oligocene suggests a close coupling between global climate and the carbon cycle, while this association appears to weaken in the Neogene (Pagani et al. 1999a , 2005). From 26 Ma to 15 Ma, warmer conditions reduced the extent of Antarctic ice, with the Antarctic ice sheet becoming re-established after that. From 10 to 5 Ma temperatures continued to drop, leading

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to the development of the marine-based West Antarctic ice sheet and ice accumulation in the Arctic. Somewhat warmer conditions prevailed in the Early Pliocene (5.33–3.2 Ma), before a trend towards the intensification of Northern Hemisphere (NH) glaciation (3.2– 2.8 Ma) (Shackleton et al. 1984, 1995a ). The onset of glaciation requires temperatures that are low enough for snow precipitation and reduced summer snow melt, and a sufficient moisture supply. The gradual shoaling of the Central American Seaway (CAS) and the establishment of the modern Atlantic-Pacific salinity contrast between 4.7 and 4.2 Ma has been considered a key event, because it led to the intensification of the North Atlantic Thermohaline Circulation (NATHC) and to increased moisture transport to high latitudes (Haug and Tiedemann 1998). Although initially this may have led to warmer conditions, the change in boundary conditions may have preconditioned the climate system towards the onset of NH glaciation (Haug et al. 2001). The final closure of the CAS around 2.7 Ma may have contributed to the crossing of a critical threshold in moisture supply and the onset of glacial conditions (Bartoli et al. 2005). In addition, Haug et al. (2005) proposed that ocean stratification in the subarctic Pacific Ocean at 2.7 Ma led to increased seasonality, with summer warmth extending into the autumn and thus providing the moisture source for snow accumulation in North America. Recent modelling experiments provide support for the notion that the closure of the CAS led to an intensification of the North Atlantic circulation and to increased precipitation over Greenland and North America, but the simulated changes in ice volume appear too small (Lunt et al. 2008). The implication is that the CAS closure does not appear to be the primary factor that caused the onset of NH glaciation and that a more likely candidate may be decreasing atmospheric CO2 concentrations (Lunt et al. 2008). This may have combined with the modulation of the amplitude of the 41kyr obliquity cycle by the longer 1.2-Myr cycle, which led to more extreme changes in tilt angle after 3.1 Ma and therefore the occurrence of cold summers in higher latitudes (Berger and Loutre 1991). According to Lourens and Hilgen (1997) the influence of the long-periodic variations in obliquity is observed during several key moments in Neogene climate history. More specifically, high-amplitude variations in tilt connected with the 1.2 Myr cycle appear to be associated with sea-level lowstands of the short-term eustatic cycles of Haq et al. (1987), implying that these cycles are glacio-eustatic (Lourens and Hilgen 1997). During the Oligocene, Wade and Pälike (2004) found that glacial events (Oi) were related to the ∼405-kyr

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0

Vegetation in the Mediterranean

Epoch Climate events Pleistocene NH glaciation Pliocene

Glacial-interglacial vegetation cycles Expansion of open vegetation

W. Antarctic ice sheet 10

Mixed temperate

Miocene Mid-Miocene climatic optimum 20

Age (Ma)

Late Oligocene warming 30

Broad leaved evergreen/deciduous

Oligocene Antarctic glaciation

Paratropical

40 Eocene Early Eocene climatic optimum

50

Tropical

60 Palaeocene

7

6

5

4

3 2 δ18Obenthic (‰)

1

0

-1

-2

Fig. 4.2. Global oxygen isotope record based on data from more than 40 DSDP and ODP sites. The curve shown is a five-point running mean. NH = Northern Hemisphere. Also shown are some global climate events and Mediterranean vegetation changes (modified from Zachos et al. 2001).

eccentricity cycle, and also to the 1.2-Myr obliquity amplitude modulation cycle. In the Miocene, glacial events (Mi) are correlated with the 1.2 Myr cycle. The Tortonian–Messinian boundary (∼7.2 Ma), which coincides with the onset of a significant sea-level drop and in the Mediterranean is characterized by a notable increase in cold planktonic foraminifera (left-coiling neogloboquadrinid species), is also associated with the 1.2 Myr cycle (Lourens and Hilgen 1997). In the PlioPleistocene, the third-order eustatic changes at around 2.8 and 1.7 Ma also appear to be related to the 1.2 Myr obliquity cycle. These events are marked in the Mediterranean marine record by the first occurrence of Neogloboquadrina atlantica (s) during MIS G10 (Lourens et al. 1996) and a strong increase of N. pachyderma (s) from MIS 64 at the Plio-Pleistocene boundary as defined in the Vrica section, southern Italy (Lourens et al. 2004). Curiously, however, whereas in the Pliocene and Pleistocene, periods of obliquity maxima, modulated

by the 1.2 Myr cycle correspond to enhanced glaciation (Lourens and Hilgen 1997), in the Miocene and Oligocene the opposite situation is observed, and low obliquity amplitude variations correspond to glacial events (Turco et al. 2001; Wade and Pälike 2004).

Global Changes: Vegetation The long-term Cenozoic trend from greenhouse to icehouse conditions had a profound influence on the history of marine and terrestrial biota. The reduction in both sea surface and air temperatures led to decreased evaporation and reduced moisture availability. In addition, cryospheric development and high-latitude cooling led to a sharpening of the meridional thermal gradient, which in turn resulted in increased wind strength, intensification of atmospheric and oceanic circulation, and near-shore cold water upwelling in certain areas

Cenozoic Climate and Vegetation Change

(Kennett 1995). Aridification was a major consequence of these changes. This was further accentuated in certain places by tectonic processes through mountain building, producing orographic rain-shadow effects, or blocking the passage of monsoonal winds. All these changes promoted the expansion of open, herbaceous vegetation. Indeed, the palaeobotanical record provides clear evidence of the emergence of plant families whose members are predominantly herbs and shrubs during the course of the Cenozoic (Singh 1988). Most of these families appeared in the Late Cretaceous/Palaeocene, Eocene, and Oligocene but did not expand until the Miocene or even later. Until the end of the Eocene the world appears to have been almost completely forested, with low and middle latitudes dominated by evergreen mega- and mesophyllous angiosperms with some gymnosperms (Bredenkamp et al. 2002). Grasses evolved sometime between 70 and 55 Ma (Kellogg 2001; Jacobs 2004), but remained restricted until the Miocene. In Africa, grasses are consistently recorded in the Early Miocene (23–16 Ma), but began to expand after 16 Ma, with widespread savanna becoming established by 8 Ma (Jacobs 2004). The history of mid-latitude grasslands (or steppes) is more uncertain, with evidence pointing to a Late Miocene appearance in Asia and North America. The first unequivocal evidence for tundra is from the Late Pliocene (Wolfe 1985 and reference therein). A major evolutionary step in the grasses was the appearance of the C4 photosynthetic pathway, which can be viewed as a CO2 -concentrating mechanism that maximizes rates of photosynthesis by eliminating the effects of photorespiration. This provides a competitive advantage over plants with the more primitive C3 pathway when the ratio of atmospheric CO2 to O2 is low (e.g. Ehleringer and Björkman 1977; Ehleringer et al. 1991). At present, the distribution of C4 grasses is strongly correlated with regions of high minimum temperatures, strong seasonal precipitation and a wet and warm growing season; C4 species of grasses today dominate the prairies of North America, the grasslands of Africa, and the llanos and cerrados of South America (e.g. Collatz et al. 1998; F. I. Woodward et al. 2004). The earliest unequivocal fossil record of a leaf that can be designated as C4 is dated at 12.5 Ma (Nambudiri et al. 1978). Stable carbon isotopic ratios have been used to distinguish C4 from C3 plants and these provide an earlier first appearance date around 15 Ma (Kingston et al. 1994). Molecular clock estimates place the origin of C4 much earlier between 25 and 32 Ma, while the phylogeny of the grass family suggests that the C4 photosynthetic pathway evolved independently several times in different subfamilies (Kellogg 2001). Between 8 and 6 Ma, a

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major expansion in C4 grasses across low latitudes has been inferred on the basis of distinct changes in ‰13 C of fossil mammal tooth enamel and palaeosol carbonates, but the causes of this expansion remain unclear. Cerling et al. (1997) have suggested that it may be related to a decrease in atmospheric CO2 concentrations below a critical level. On the other hand, Pagani et al. (1999b), using alkenone-based CO2 reconstructions, showed that atmospheric CO2 concentrations had already stabilized at pre-industrial levels of ∼290 ppmv before the C4 expansion. Thus, CO2 changes could not have been the forcing mechanism behind this ecological shift, which according to Pagani et al. (ibid.) was driven by the enhanced low latitude aridity or changes in seasonal precipitation patterns. If that were true, however, then separate C4 expansions should have happened several times earlier at relatively more arid locations (Cerling et al. 1997). It is of course possible that local earlier expansions did take place but have not yet been documented. It is also possible that instead of their Miocene expansion, it was the initial Cenozoic evolution of the C4 photosynthetic pathway between 25 and 32 Ma that was driven by a drop in atmospheric CO2 levels. Indeed recent reconstructions of atmospheric CO2 , spanning the entire interval 5–45 Ma, show that concentrations fell from ∼1,500 to 1,000 ppmv around 35 Ma to preindustrial levels by 25 Ma (Pagani et al. 2005). The level of 500 ppmv below which the C4 photosynthetic pathway is favoured over C3 was first breached c.30 Ma and this confluence may suggest a causal relationship (Pagani et al. 2005), although tighter chronological control on the timing of the C4 origin is needed to test this idea further. The Late Miocene expansion of C4 grasslands, on the other hand, has recently been attributed to a fundamental increase in fire regimes as a result of the onset of marked seasonality characterized by a warm, moist growing season with high biomass production, followed by a dry season that would convert the biomass into highly combustible fuel (Keeley and Rundel 2005). An important point that emerges is that the establishment of a frequent fire regime was a key factor not only for the expansion, but also for the maintenance of C4 grasslands to the present day (Keeley and Rundel 2005). This is supported by modelling experiments, showing that in the absence of fire, areas that are today dominated by C4 grasslands and savannas have the climatic potential to form forests (Bond et al. 2005).

The Mediterranean World The Cenozoic environmental changes outlined above occurred against a background of continuous

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geodynamic evolution in the Mediterranean region. Jurassic geological reconstructions indicate the presence of a wedge-shaped equatorial ocean, named Tethys, separating northern from southern continents (Smith et al. 1992). The opening of the Atlantic Ocean after 170 Ma led to the collision of Africa and Eurasia and the gradual elimination of the Tethys Ocean. Today, none of the former Early Jurassic or older Tethys ocean floor remains, with the western Mediterranean underlain by Miocene and younger crust, while the eastern Mediterranean is floored by mid-Jurassic or younger crust (Livermore and Smith 1985). The connection to the Indian Ocean was maintained until the Middle Miocene; the timing of the final closure of this eastern gateway is not well-constrained, with different studies providing estimates ranging from 16 to 11 Ma (e.g. Dercourt et al. 1993; Yilmaz 1993; Jacobs et al. 1996). The collision between Africa and Eurasia gave rise to new mountain chains (Chapter 1) and also led to the separation of the Mediterranean from an eastern European epicontinental sea, the Paratethys (whose remnants today are the Black, Caspian, and Aral Seas). The separate evolution of the Paratethys started as early as the Late Rupelian (30–28 Ma) (Dercourt et al. 1993), but was completed by the Early Tortonian (∼11.6 Ma) (e.g. Sprovieri et al. 2003). This meant that the Paratethys no longer received marine waters, while the Mediterranean was deprived of the fresh waters of the eastern European rivers, which emptied into the Paratethys (Hsü et al. 1977). In the western Mediterranean, the connection to the Atlantic was maintained through two gateways, the Betic Corridor through southern Spain and the Riffian Corridor through northern Morocco (Krijgsman 2002). Arguably the most dramatic episode in the history of the Mediterranean Sea is the Messinian Salinity Crisis (MSC) some time between 6 and 5 Ma, which refers to its isolation from the Atlantic Ocean and its partial desiccation. During this time, large saline lakes replaced marine basins while some areas dried out completely, leading to the emplacement of huge thicknesses of evaporitic deposits (e.g. Hsü et al. 1977; McKenzie et al. 1999). The causes of this isolation (glacio-eustatic vs tectonic) have been much debated (Chapter 1), but the controversy appears to have been resolved by the development of an astronomically calibrated chronology based on tuning of sedimentary cycles seen across the Mediterranean to orbital changes (Krijgsman et al. 1999). This shows that the beginning and end of the MSC occurred at 5.96 Ma and 5.33 Ma, respectively, and that these changes were synchronous throughout the Mediterranean. Moreover, this precise

chronology allowed the placing of these events within the benthic ‰18 O framework (e.g. Shackleton 1995a). This showed that the onset of the MSC was not related to prominent glacio-eustatic sea-level falls of either glacial stages TG22 or TG20 (Figure 4.3), and pointed to a predominantly tectonic explanation, although orbitally driven long-term changes in sea level may have played a part (Krijgsman et al. 1999). The end of the MSC may have been associated with sea-level rise of interglacial stage TG5 (McKenzie et al. 1999), although Vidal et al. (2002) point out that this was not as prominent a glacio-eustatic event as that of TG9 (Figure 4.3) and suggest that the end of the MSC was also not directly controlled by sea-level changes. Be that as it may, the so-called ‘terminal flood’ of 5.33 Ma reestablished the connection to the Atlantic through the Strait of Gibraltar and marks the onset of the Pliocene. During the MSC, the substantially reduced water body would have led to reduced moisture availability and aridification. It has also been suggested that the absence of a Mediterranean salty outflow to the North Atlantic may have led to a weakening of deep water formation and promoted cooler conditions (Vidal et al. 2002). In biogeographical terms, the desiccation permitted animal migration across land masses, while the terminal flood led to the isolation of animals on Mediterranean islands and subsequent evolutionary changes (Azzaroli and Guazzone 1979/80). In summary, the present Mediterranean can be viewed ‘as the most recent of a series of “Mediterraneans” whose shape and area have evolved rapidly throughout Mesozoic and Cenozoic time’ (Livermore and Smith 1985: 86). The familiar configuration of the basin is a geologically recent development, with the isolation of the Mediterranean as an inland sea occurring in the Middle Miocene and the present outline of land and sea emerging around 5 Ma. With regard to the major changes in vegetation in the Mediterranean during the course of the Tertiary, a synopsis can be provided through global vegetation reconstructions and overviews by Wolfe (1985), Janis (1993), and Willis and McElwain (2002). In the Early Eocene climatic optimum (∼50 Ma), the precursors of the Mediterranean lands were covered by a tropical rainforest, which extended to 50◦ N (Figure 4.4a). These evergreen and laurel forests are considered to have been derived from the ‘Palaeotropical Geoflora’ (or ‘Tethys flora’) and represent an important component of Tertiary vegetation in Europe (Mai 1989). A ‘paratropical’ rainforest (sensu Wolfe 1985) extended between 50◦ and 60–5◦ N, and up to 70◦ N in coastal areas; this contained a mixture of tropical and temperate elements with mangrove

Cenozoic Climate and Vegetation Change

95

TG5

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‰ 2

2.5

3

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4 4.0

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5.8

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6.2

6.4

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Age (Ma) Fig. 4.3. A section of the compiled oxygen isotope record of Figure 4.2 (Zachos et al. 2001) for the interval 4–7 Ma. The duration of the Messinian Salinity Crisis is indicated (see Chapter 1), as are some of the prominent Marine Isotopic Stages (e.g. TG5) for that period.

swamps along the coasts and is classified under the summer-wet biome (Willis and McElwain 2002). In the continental interiors, a broad-leaved evergreen forest occurred between 60–5◦ N and 70◦ N, and finally north of 70◦ N formed a polar vegetation dominated by broadleaved deciduous trees and deciduous conifers (Wolfe 1985). During the cooling trend of the second half of the Eocene (Figure 4.4b), tropical rainforest in the Mediterranean region was gradually replaced by paratropical rainforest; broad-leaved evergreen forest migrated south of 60◦ N, while to the north of that extended a diverse mixed coniferous forest (Wolfe 1985). Following the Oligocene initiation of the East Antarctic ice sheet and further global cooling, vegetation in the Mediterranean region became increasingly dominated by broad-leaved deciduous forests, which extended over large areas of the NH (Figure 4.4c). The majority of the deciduous component represents Arctotertiary species, which had already begun entering Europe in the Palaeocene/Eocene, but during the course of the Oligocene invaded all European forest communities, producing very rich mixedmesophytic forests (Mai 1989). Broad-leaved evergreen vegetation gradually retreated south of 30◦ N (Wolfe 1985; Janis 1992) and became increasingly temperate in character (Mai 1989). The warmer conditions from 26 Ma to 15 Ma, culminating in the Mid-Miocene climatic optimum, led to a northward expansion of broad-

leaved evergreen forest up to 45◦ along the coasts (Wolfe 1985), but the ensuing steady decline in temperatures led to the near complete extinction of European laurophyll species (Mai 1989) and the establishment of mixed temperate woodlands (Quade et al. 1994 and references therein). The decline in temperature, the elimination of the Tethys Ocean and mountain building all contributed to the aridification of Africa and the expansion of open vegetation (Singh 1988). With regard to the state of vegetation during the MSC, palynological studies from the Tyrrhenian basin show the presence of deserts and saline terrain vegetation interspersed with riparian or delta forests in the former abyssal plain and lowlands; a thermophilous mid-altitude zone with subtropical and temperate elements and a montane zone are reconstructed further inland (see review by BertolaniMarchetti 1985). Of particular significance are palynological results from the Ptolemais and Servia basins in northern Greece, spanning three phases: before (6.75– 6.7 Ma and 6.33–6.28 Ma), during (5.44–5.21 Ma), and after (4.36–4.15 Ma) the MSC (Kloosterboer-van Hoeve 2000). Because these results come from the same area they can be used to assess directly the character of these periods without any complications arising from differences in local factors when attempting long-distance comparisons. Vegetation in the pre-MSC phase was characterized by a mixture of deciduous and coniferous

(a) 120°W

150°W

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150°E

120°E

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90°E

60°W

60°E

30°W

0°

30°E

(b)

Mixed Coniferous Broad-leaved Deciduous Broad-leaved Evergreen Paratropical Tropical

(c)

Fig. 4.4. Northern Hemisphere palaeogeography and global vegetation maps for selected time slices in the Tertiary: (a) Early Eocene (∼50 Ma), (b) Middle/Late Eocene (∼37 Ma), and (c) Late Oligocene (∼27 Ma). The dashed line within the Broad-leaved Deciduous forest in Figure 4.4c indicates the border between Mixed Northern Hardwood forest in the north and the more southern Broad-leaved Deciduous forest. Modified from Wolfe (1985), with additional information from Janis (1992).

Cenozoic Climate and Vegetation Change

trees with an almost complete absence of herbs. During the MSC, the record is strikingly different, showing strong cyclic (precessional) changes between periods of herbaceous dominance and coniferous dominance; transitional intervals show increases in deciduous and evergreen Quercus. In the post-MSC phase, the overall abundance of herbs decreased (but not entirely) and periods of evergreen and deciduous Quercus woodland alternated with pine-dominated woodland. The overall impression is of subdued cyclicity with moist forest conditions before the MSC, prominent aridity and seasonality during the MSC, and the establishment of more modern open woodland following the MSC, but with seasonality persisting. Considerably more information is available for the Pliocene, providing an insight into the regional differentiation around the Mediterranean Basin (the reader is referred to reviews by Axelrod 1973 and Suc et al. 1995a, b). For the Early Pliocene (Zanclean (5.33–3.6 Ma)) the vegetation reconstructions have been organized into the following geographical quadrants. The north-western region (from Catalonia to central Italy) was characterized by a strong subtropical element. Taxodium swamps were found in coastal plains, mid-altitude forest belts were occupied by Sequoia, Cathaya, Cedrus, and Tsuga, and higher-altitude forest belts by Abies and Picea (Suc et al. 1995a, b). The south-western region was dominated by herbaceous xerophytic vegetation (Asteraceae, Plantago), including some subdesertic elements. Mediterranean plants were regularly represented and some tropical plants were present in the southernmost parts (ibid.). In the south-central region, evidence from Calabria and Sicily shows a gradually decreasing subtropical component, and the steady presence of mediterranean elements and herbaceous vegetation (Bertoldi et al. 1989). The north-eastern region had temperate forests alternating with pine-dominated forests or herbaceous vegetation (Drivaliari 1993; Kloosterboer-van Hoeve 2000). Finally, in the south-eastern region the Nile area was dominated by open vegetation, which included desertic elements (Drivaliari 1993). In the coastal plain of Israel, a mediterranean vegetation was established along with Artemisia (Horowitz 1974). An intriguing pollen record of Pliocene age from the Plateau Tufa, Kurkur Oasis (west of the Aswan Dam) shows the presence of taxa of the Arctotertiary Geoflora (Alnus, Corylus, Salix, Ostrya, Betula, Tilia, Aesculus, Carpinus, Quercus, Platanus, Ulmus) along with mediterranean (e.g. Oleaceae, Rhamnaceae, Cistaceae, Pistacia, Celtis) and subtropical African (e.g. Podocarpus, Phoenix, Loranthoideae) elements; the record also shows an increase

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in halophytic and xerophytic elements (van Campo et al. 1968). The Plateau Tufa record is reminiscent of another record of Early Pleistocene age from the Hoggar Mountains (1,830m a.s.l.) in the central Sahara where pollen of Tilia, Quercus, Alnus, Juglans, Pterocarya, Zelkova, Ulmus, Picea, Pinus, Ostrya, Corylus, and Fraxinus were found along with mediterranean and tropical elements (van Campo et al. 1964). Together the two records lend support to the notion that elements of the Arctotertiary Geoflora had penetrated into the central Sahara at the end of the Tertiary (Axelrod 1973). Pollen records from the early Late Pliocene (Piacenzian, 3.60–2.59 Ma) show a gradual reduction of the subtropical component and an increase in herbaceous vegetation. The initiation of major NH glaciation is marked in the Mediterranean region by the expansion of steppe (Suc 1984) and the start of glacial–interglacial vegetation cycles. After the PlioPleistocene boundary (1.81 Ma) there is a further increase in the abundance of herbaceous vegetation during glacial intervals in the Mediterranean (Combourieu Nebout and Vergnaud-Grazzini 1991). In parallel with the intensification of glaciation during the course of the Late Pliocene and Pleistocene, pollen records show the progressive disappearance of subtropical taxa from the Mediterranean and Europe. Several species disappeared right at the onset of major NH glaciation, while others persisted but at much reduced abundances. Over the next ∼2 Myr there was a continuous extinction of tree species from Europe. By about 600–400 ka (depending on location), the composition of European forests was similar to the present-day situation. The causes for these extirpations will be discussed in detail in the macro-scale section. In floristic terms, the present-day sclerophyll elements in the Mediterranean are considered to have their origins in the Tertiary laurophyll vegetation (Mai 1989). Palamarev (1989) provides a survey of the fossil occurrence of plant species of the ancient Tethyan and Paratethyan communities, which developed into the Mediterranean forest vegetation by the end of the Tertiary (Table 4.2). According to this, a small number of species appeared first in the Eocene, while twenty-three species have an Oligocene first appearance. The Miocene list of taxa is the most diverse with forty-nine new species, while five new species make their first appearance in the Pliocene fossil record. These records show that the major part of palaeomediterranean woody species appeared in the Miocene mixed evergreen and deciduous forests (Palamarev 1989). The Late Tertiary sclerophyllous

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Chronis Tzedakis TABLE 4.2. List of genera of species’ first fossil appearance in the Tertiary record (from Palamarev 1989)

Eocene Oligocene Miocene

Pliocene

Pinus, Tetraclinis, Periploca, Platanus, Rhamnus, Chamaerops, Smilax Abies, Cupressus, Picea, Pinus, Acer, Arbutus, Carpinus, Celtis, Ceratonia, Cercis, Coriaria, Cotinus, Laurus, Myrtus, Olea, Ostrya, Pistacia, Punica, Quercus(evergreen), Viburnum, Smilax Abies, Cedrus, Juniperus, Pinus, Acer, Alnus, Anagyris, Buxus, Cerasus, Colutea, Daphne, Ficus, Fraxinus, Lonicera, Marsdenia, Nerium, Paliurus, Phillyrea, Pistacia, Prunus, Pyracantha, Quercus (including semi-evergreen and deciduous), Rhus, Styrax, Tamarix, Viburnum, Vitex, Ruscus Ephedra, Juniperus, Ilex, Jasminum, Prunus

forests also contained deciduous oaks, which appeared in the subtropical flora of southern Europe in the Early Miocene as montane elements and then descended to lower belts (Palamarev 1989). While the sclerophyllous floral elements appeared early in the Cenozoic, they only became widely established in the Mediterranean after the disappearance of the laurophyll vegetation (Mai 1989). In terms of the history of C4 grasses in the Mediterranean, stable isotopic evidence from palaeosol carbonates and fossil teeth in Greece show that Miocene and Pliocene vegetation was dominated by C3 plants, mainly representing forest and woodland with a small herbaceous component, with no evidence for the presence of C4 grasses (Quade et al. 1994). The present-day distribution of C4 plants in the Mediterranean is limited, presumably a reflection of the absence of a wet growing season: modern maps show minimal abundance in most places with somewhat elevated presence (100

Fig. 6.3. Comparison of soil depths as measured in different climate zones and on different lithologies on the island of Lesvos (modified from Kosmas et al. 2000).

Mediterranean Soils Distribution and Types of Soil Yaalon (1997) has provided an overview of Mediterranean soils, considering their classification and genesis. While earlier soil-classification systems had used terms such as red and brown Mediterranean soils, this use largely fell out of practice from the later 1960s. In the revised FAO classification (see below), these soils are now considered to be luvisols or cambisols. A major element of the debate was the common perception of the dominance of red soils, which Yaalon considers to be related to the oxidation of iron oxides (Figure 6.4) under dry, summer conditions. As such, red soils from this source would have to be relatively recent, or relate to past conditions of similar conditions (e.g. Woodward et al. 1994), based on our knowledge of climate conditions. Another mechanism of reddening would be heating during wildfires, which can also bring about similar transformations to haematite at temperatures between 200 and 400◦ C. Recent work by Wondafrash et al. (2005) has suggested that relatively permanent changes in colour of the surface horizons of Mediterranean soils can come about through burning. In this case, the persistence of red-coloured horizons at depth in the soil will depend on the relative dissolution and erosion rates and fire frequency. While red soils—commonly using the term terra rossa—are generally associated with limestone bedrock,

it is now appreciated that red soils also occur on other lithologies, including igneous, metamorphic, and conglomerates (Yassoglou et al. 1997). MacLeod (1980) pointed out the problem of assuming the reddening of limestone-derived soils could come from in situ alteration of non-carbonate residues, given that the latter are typically very small (8 3–x

Fe (OH)x Precipitation

3–g

Fe (OH)g

Rapid oxidation

Oxidation, Deprotonation

Fe2+

Fig. 6.4. Processes leading to the formation of iron oxides and thus the development of brown (goethite) and red (haematite) soils. Thicker arrows are the more common processes (modified from Yaalon 1997).

Tomadin et al. 1984, 1990) and north-western Mediterranean (Bücher 1989; Durand et al. 1992) have also been noted. Loess soils, where dust deposition exceeds about 25 mm ka−1 , can be found in Israel (Issar et al. 1989), Tunisia, and Morocco (Thomas 1997). However, measurements of present-day dust accumulation suggest that the loess is not actively forming, but again relates to past fluctuations in the position of the ITCZ. Figure 6.5 shows the spatial distribution of soils in the Mediterranean according to the FAO classification. While there is again a north–south distinction in the

types of soil present, the pattern (even as greatly simplified in this map) is more complicated than that of the climate-driven process map (Figure 6.1), providing further support for the Pope et al. (1995) model of local conditions controlling the type of soil-forming process. In the north, cambisols are the most frequent type (Table 6.1), dominated by eutric, calcic, and dystric cambisols in almost equal measure. Luvisols are next in importance, with chromic, orthic, and then gleyic luvisols being the main types present. Lithosols are the next most frequent type, followed by fluvisols (mainly calcaric) and xerosols (again mainly calcic). No other soil type covers more than 4 per cent of the total area. In the south, yermosols (dominantly calcic followed by undifferentiated with a small proportion of gypsic yermosols) are most frequent, followed by lithosols and podzoluvisols. Xerosols (mainly calcic) are around twice as frequent as (mainly calcic) fluvisols—mostly in the Nile valley—and cambisols. The latter are less than ten times as frequent in the south as in the north. These soil types make up nearly 90 per cent of the total area in the south. It should be clear that Yaalon’s (1997) description of ‘typical’ Mediterranean soils relates only to the northern part of the basin, and even so, the notion of a typical soil conceals a great deal of variability. Figure 6.6 illustrates a number of example soil profiles from the northern part of the basin. It should be no surprise that soils in the south are poorly developed, relating to the arid regime (almost 50% are yermosols or xerosols, which are characterized by a weakly developed ochric A horizon and a further 17% are lithosols, which are characterized by having a hard rock level within 10 cm of the surface) and the higher prevalence of wind erosion (Chapter 14). Comparison of the climate-based characterization and the FAO soil classification shows some convergence—for instance the majority of cambisols and luvisols occur in the moderate chemical weathering area, and most yermosols are in the very slight weathering zone—but also some differences. Nearly two-thirds of the lithosols are in the very slight weathering area, with just over a third in the moderate chemical weathering class; very few occur in zones characterized climatically as being dominated by mechanical weathering. While this pattern may in part be due to the scale of mapping and inadequacies with the climate model, it may also suggest that these soils are relict from past climate regimes. Alternatively, their location, which is dominantly in an arc through the Balkans and into Asia Minor and southern Turkey may reflect areas where high winter precipitation biases the climate model. Almost as many xerosols occur in areas mapped as undergoing moderate chemical

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Fig. 6.5. Spatial distribution of soils in the Mediterranean basin according to the FAO classification scheme (modified from Wainwright and Thornes 2003, based on data provided by NOAA 1984).

TABLE 6.1. Comparison of proportions of different soil types according to the FAO classification relative to location in the north or south of the Mediterranean basin North Cambisols Luvisols Lithosols Fluvisols Xerosols Rendzinas Chernozems Kastanozems Podzols Phaeozems Regosols Vertisols Rankers Planosols Andosols Gleysols Acrisols Solonchaks Podzoluvisols

%

South

%

42.1 16.4 8.9 4.6 4.3 3.5 3.5 3.1 2.4 2.3 2.1 2.0 1.0 0.8 0.7 0.6 0.6 0.3 0.3

Yermosols Lithosols Podzoluvisols Xerosols Fluvisols Cambisols Luvisols Solonchaks Regosols Rock Rendzinas Kastanozems Vertisols Salt

41.2 16.7 15.2 8.1 4.3 3.9 2.6 2.2 1.8 1.0 1.0 0.9 0.5 0.2

Source: Analysis is based on the NOAA-AVHRR (1984) data (after Wainwright and Thornes 2003).

weathering as being subject to very slight weathering. These possible anomalies are located in the northern Atlas and in central Anatolia. The idea that soil-forming processes have changed significantly through time has already been considered in terms of controlling mechanisms. Further lines of support for this idea come from the study of palaeosols. For example, Leone et al. (2000) have described a 150-m stratigraphic section in central Italy that is interpreted to have been deposited over a period of 100 to 300 ka in the Late Pliocene, which contains seventy-three palaeosols as well as lignite layers. Pollen evidence shows the presence of subtropical forest species, although isotopic analyses suggest that precipitation was not much higher than at present, and also points to the presence of grass vegetation. Günster and Skowronek (2001) have investigated sequences in the Granada basin of southwest Spain that contain sixty-five palaeosols within a thickness of up to 70 m of alluvial fan sediments. Early Pliocene and Early Pleistocene palaeosols show higher redness values, suggestive of more subtropical conditions, while Later Pliocene soils are more similar to those

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a Ah BT1 BT2 C/R

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Fig. 6.6. Examples of typical soil profiles from the Mediterranean, as described by Tavernier (1995): (a) terra rossa on a cherty limestone near Rieti in Italy; (b) a pellic vertisol derived from igneous rocks in southern Portugal; (c) a ferric luvisol derived from schists in Portugal; (d) a ‘Black Mediterranean Soil’ (pellic vertisol) derived from marls near Florence in Italy; (e) a calcaro-pellic vertisol derived from Tertiary, calcareous lacustrine deposits in Thrace, Greece; (f) a calcaro-vertic cambisol derived from lagoonal clays in south-east Italy; (g) a calcivertic luvisol derived from Miocene clays near Badajoz, Spain; and (h) a calcic xerosol derived from Quaternary alluvium in Murcia, Spain (modified from Wainwright and Thornes 2003).

actively forming and contain up to sixteen pedogenic calcretes, which would also suggest at least seasonal aridity. In the Pleistocene, Ortiz et al. (2002) investigated soils on relatively stable alluvial fan surfaces in this same basin. They found that soils formed up to about 240 ka had higher amounts of clay and iron oxides, suggesting greater pedogenic activity during warmer periods. In some of the older soils, micromorphological evidence of frost shattering suggested mechanical weathering that the authors attributed to activity in cold periods (a feature also noted by Günster and Skowronek 2001, in mid–Late Pleistocene palaeosols in the area). Furthermore, older soils had greater CaCO3 accumulations in the C horizon, interpreted as being due to infiltration with carbonate-rich runoff, although input of carbonate from dust cannot be excluded. Indeed, Pulido-Villena et al. (2006) have measured average inputs of 1.6 g m−2 a−1 of Ca in high mountain lakes in Spain. Late Pleistocene–Early Holocene palaeosols in Israel (Gvirtzman and Wieder 2001) have been related to more humid phases relating to ITCZ fluctuations as discussed above. In the Pyrenean scree slopes already discussed, García-Ruiz et al. (2001) ascribe a phase of carbonate cementation to warmer, more humid conditions in the early Holocene. Van Andel and

Runnels (1987; van Andel et al. 1990) also describe early Holocene palaeosols from different areas in southern Greece.

Soils as a Resource As noted by Sumner and Wilding (2000: p. xvii): ‘We owe our existence to the extremely thin but precious skin, called soil, which covers the unweathered and partially weathered geological formations at the Earth’s surface with a unique and extremely thin, fragile layer.’ Soils store water and provide the physical and nutrient support for vegetation (Chapter 7) and thus animals (Chapter 5). As has been already noted, where older soils are not preserved, Mediterranean soils tend to be rather thin—often only a few tens of centimetres thick, and some slopes are bare, bedrock surfaces devoid of any soil cover. One consequence of thin soils is that they have a relatively low water-holding capacity, which will tend to promote runoff and erosion (see below), which provides a positive feedback to maintain relatively thin soils. This feedback operates not only by the physical removal of sediment, but also by the removal of nutrients, which, combined with at least seasonal aridity, maintains lower vegetation cover that also tends to promote erosion (see discussion in Thornes 1985, 1990). However, as

Weathering, Soils, and Slope Processes

also discussed, Mediterranean soils are characterized by variability and it would be incorrect to state that all soils are poor. Based on the FAO classification, 16 per cent of soils in the northern basin are classified as having a good nutrient status (mainly cambisols, but also some regosols and fluvisols), although not surprisingly, the value is only 1.3 per cent in the south. However, deep soils exist throughout the basin (approximately 4% of both the north and the south have fluvisols), in part as a response to deposition in areas of low relief of sediment eroded from uplands (where it is perhaps less useful as a resource), and a positive effect of the stoniness of soils is the maintenance of more humid soil microclimates (Casals et al. 2000), with stones acting as a mulch during periods of dryness (van Wesemael et al. 2000; Larcheveque et al. 2005). During the Late Pleistocene and through the Holocene, there has been increasing human impact on soils (Wainwright and Thornes 2003). The impacts of human activity on erosion will be discussed below, but it is important to recognize other negative and positive changes. Modification of the form of hillslopes through terracing and deliberate movement of soils (e.g. Gams et al. 1993; see also Figure 10.12 in Chapter 10) have been commonplace for millennia. These changes to the surface have the effect not only of increasing water-storage capacity and thus vegetation, but also tend to reduce the slope angle and thus minimize erosion (Van Andel and Runnels 1987). Traditional cultivation practices have often (but not always) maintained a more effective use of limited soils, and integrated agropastoral systems provide an effective means of recycling nutrients. Pollution of soils is also not restricted to the modern period. For example, Aberg et al. (2001) show evidence of soil contamination from lead smelting in the Hellenistic and Roman periods in Greece. It is clear that the pollution of soils has accelerated greatly since the mechanization of agriculture from the 1950s (e.g. Zalidis et al. 2002). Salinization is also a widespread (and long-lasting) problem, albeit often more accentuated in the southern and eastern parts of the basin. As elsewhere, pollution can also be related to distant sources. For example, Vavliakis et al. (1990) suggested that dissolution rates of dolomitic marbles in northern Greece doubled when dominant winds from eastern and central Europe increased the acidity of local rainfall to pH 90 per cent of its length, 85 per cent of its depth, 60 per cent of its area, and 35 per cent of its volume in the first 5 per cent of its lifetime. The temporal scaling of erosion rates is thus clearly non-linear. This effect can be seen further in the study of Wise et al. (1982), who argued on archaeological grounds that some badland gully systems in southern Spain had evolved little since the Bronze Age, and thus the gullies could be responsible for very little current erosion. Over shorter timescales, there are also important differences—at the same Portuguese and Spanish sites mentioned above, gullies were only responsible for 47–51 per cent of total erosion during a single wet year, compared to the much higher values cited above, which are averages over 3–20 years (Poesen et al. 2002). In extreme cases, especially in highly erodible soils, gullying can expand dramatically, producing entire landscapes that are dominated by these features. These badlands are extensively found in southern Spain, both on marls and unconsolidated bedrocks (Figure 6.15), in south-eastern France in the Black Marls of the Alps, in Italy on marine clays, and in northern Greece (see discussion in Chapter 1). Essentially, badlands relate to the extension of gullies until their contributing areas are reduced to be unable to maintain concentrated overland flows. There are important feedbacks to vegetation growth, with the highly eroded substrates of the badlands being unable to support vegetation due to the low water-holding capacity, high runoff rates, and low nutrient availability. Subsurface erosion can also be an important phenomenon in the Mediterranean, especially in areas of unconsolidated sediments. Soil pipes may occur just beneath the surface, especially if sediments remain disaggregated beneath a crusted surface, allowing them to be detached by vertical and lateral subsurface flows. Piping can also arise due to dissolution, for example in soils or sediments where gypsum or calcium carbonates are present, and may also occur at depth (Harvey 1982). Soil macropores (due to decayed roots or animal burrows, for example) or deep desiccation cracks can allow water to infiltrate rapidly to depth, eroding the macropore or crack as it moves. Such flows may also concentrate along lithological boundaries or structural fissures. Tunnel systems can typically be very

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Fig. 6.13. Aerial photograph of Stipa tenacissima vegetated slopes at the Rambla Honda in southern Spain, showing different patterns of connectivity between bare patches (light coloured) in different parts of the catchment, resulting in significant increases in erosion rates in areas where bare patches are interconnected (see Puigdefabregas et al. 1999). Image courtesy of Juan Puigdefabregas.

extensive in susceptible areas (Figure 6.16) and be responsible for rapid transfers of water and sediment into the channel system, often even before main flows across the surface have begun (Farifteh and Soeters 1999). Continued pipe erosion tends to be unstable, leading to subsidence and collapse of the ground surface. Gutierrez-Santolalla et al. (2005) demonstrate the effects of extensive pipes and consequent hazards from gypsum dissolution in the Ebro basin. Pipe collapse is a common cause of rill initiation—Faulkner et al. (2004) discuss the process in dispersive soils in southern Spain—as well as gully initiation (Haigh 1990; Imeson et al. 1982; Martin Penela 1994), not least because collapse can be related to the presence of significant amounts of flowing water under turbulent conditions (Figure 6.17).

While some effort to estimate the relative importance of different water-erosion processes is made above, this chapter does not present a summary table of different measured rates (q.v. Poesen and Hooke 1997 and Table 6.2 in Wainwright and Thornes 2003). The reason for this omission is due to problems that occur when trying to compare erosion rates that are measured at different spatial scales. Parsons et al. (2004) discuss the spatial scaling of erosion rates in detail and show that rates that are measured as specific yields (or areal averages, i.e. in units of M L−2 T−1 ) produce meaningless results. For example, the median rate of the erosion rates in Table 6.2 of Wainwright and Thornes (2003) when expressed in terms of ground lowering is 69.3 mm ka−1 . This rate is in excess of most of the soil-production rates

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Fig. 6.14. Rills formed during the extreme rainfalls of 22 September 1992 in south-eastern France (photo: John Wainwright).

as discussed above, and implies that rather than there being thin soils throughout the Mediterranean, there should in fact be none at all! Although it can be argued that high values in the data relate to human activity,

rates reported under evergreen oak forest in north-east Spain (Sala and Calvo 1990) equate to ground-lowering rates of 156.31 mm ka−1 , while Diamantopoulos (1993) reports rates equivalent to between 19.25 mm ka−1 and

Fig. 6.15. Badlands at Tabernas, Almería, Spain (photo: John Wainwright).

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Pipes Drainage network

N 0

500 metres

Agri Valley

Fig. 6.16. Map of observed pipes showing the extent of subsurface erosion in the Agri basin, southern Italy (modified from Farifteh and Soeters 1999).

167.86 mm ka−1 under maquis vegetation in Greece, so this argument does not necessarily always hold. Parsons et al. (2004) demonstrate that the fallacy with the use of specific yields is the assumption that sediment, once detached, remains in transport; discussion of transport distances above suggests that most sediment, especially in unconcentrated flows, travels only a very short distance. Specific yields thus drastically overestimate movement because of redeposition within slope units. The detailed analysis of Martínez Casasnovas

et al. (2002) shows this point very clearly, with extensive areas of redeposition mapped even during an extreme storm event (Figure 6.18).

Processes of Slope Failure and Mass Movement Slope failure occurs when the gravitational force acting in a downslope direction on a mass of soil or rock exceeds the resistance of that mass to movement. The

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Fig. 6.17. Linked pipe and gully erosion, Murcía, Spain (photo: John Wainwright).

corresponding mass movements cover a wide range of styles and size of movement of material. Shallow translational slides or rock topples may involve less than a cubic metre to several cubic metres of material. At the other end of the extreme, large rock avalanches and slides or mudflows may involve even cubic kilometres of material (Figure 6.19). Characterization and definition of different types of movement are described in detail by Dikau et al. (1996). Guzzetti et al. (2002) have investigated the pattern of landslides in relation to the size of movement in areas of central Italy and demonstrated that distributions of landslides triggered by different mechanisms or by single events seem to follow similar patterns of area-frequency relationships (Figure 6.20), which they relate to models of selforganized criticality. Mass movements are a very common process in the Mediterranean and submarine mass movements have also been widely reported (Lasras et al. 2004; Chapter 1). van Asch (1986) suggests that the terrestrial environment is particularly prone to the phenomenon because of the lithology of surface materials, recent and ongoing

seismic activity, climate conditions, and the nature of vegetation cover, all accentuated by the general steepness of slopes. Lithologically, clays commonly occur in soils and weakly consolidated bedrock, as discussed above. Recent sediments are often weakly, if at all, lithified, which further reduces their strength. Elsewhere, tectonic activity has produced significant amounts of fractures and faults, which can act as planes of weakness (Mariolakis 1991). Overfolding in the high mountains often produces the emplacement of rocks above marls and clays, which leads to the possibility of large slides. These areas also contain relict periglacial and glacial deposits above impermeable horizons, further accentuating movement. Uplift also leads in general to oversteepening of slopes (especially when coupled with the removal of sediments at the base by high energy river channels) and thus failure (e.g. Thornes and Alcántara Ayala 1998) (Figure 6.21). Seismic activity also leads to the accentuation of the gravitational force on slopes, and is responsible for a significant number of landslides in the region. Shallow landslides tend to occur as a result of high intensity rainfall events (e.g.

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Hillside ditch Vine row

Altitude differences (m)

-0.4 – -0.3 -0.3 – -0.2 –0.2 – –0.1

erosion

–0.1 – 0 0 – 0.1 0.1 – 0.2 0.2 – 0.3

deposition

0.3 – 0.4 0.4 – 0.5

0

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50 metres

Fig. 6.18. Map of erosion (negative altitude differences) and deposition (positive altitude differences) following an extreme storm event of 215 mm (205 mm falling in 2.25 h, with an average intensity of 91.8 mm h−1 , and a maximum 30-min intensity of 170 mm h−1 ) in Catalonia, Spain in June 2000 (modified from Martínez Casasnovas et al. 2002). The detailed patterns of redeposition within the field demonstrate problems with standard approaches of measuring erosion rates as specific yields (see Parsons et al. 2004).

Wainwright 1996c), but longer periods of rainfall— typically in the winter months—can trigger much deeper-seated failures (e.g. Capecchi and Focardi 1988; Flageollet et al. 1999; Maquaire et al. 2003). Snowmelt is also important at higher altitudes—as seen by the 4,233 landslides mapped by Guzzetti et al. (2002) as a result of a single rapid snowmelt event in Umbria in January 1997 (Figure 6.20). The sparseness of vegetation cover on slopes can also increase the risk of shallow failures, especially on agricultural or sparse matorral areas (van Asch 1980), although the factor is irrelevant for the larger failures, where roots have little or no effect on cohesion of the surface. In some cases, dense vegetation can also lead to failure, by overloading slopes. There is also a possible positive feedback in the occurrence of landslides. Once activated, slope materials can become weakened, and thus slope failure more easily recurs in the same location. Maquaire et al. (2003) investigated several large earthflows in the Barcelonette basin in the French Alps, and suggested that once triggered, surface materials would be likely to remain weakened for up to

150 years (Figure 6.22). This weakening would also produce feedbacks to other processes, such as the lowering of thresholds for the initiation of gullying. Mass movements are responsible for a significant amount of sediment mobilization in the region (e.g. Keefer 1994). Ergenzinger (1992) carried out a detailed analysis of frequency–magnitude relationships for a basin in Calabria and showed that fluvial activity in this basin is controlled by inputs from major, earthquakeinduced landslide events. It can be argued that major landslides are, at least in the longer term, likely to be the most important landscape-modifying process in the Mediterranean landscape.

Slope Hazards Notwithstanding the caveats mentioned above about the extrapolation of soil-erosion rates from small-scale data, it is inescapable that the Mediterranean region has undergone significant periods of soil loss in the past and

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Fig. 6.19. Examples of different types of mass movement: (a) shallow translational slide, French Pre-Alps (photo: John Wainwright); and (b) Super-Sauze earthflow, French Alps (photo: Olivier Maquaire).

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Fig. 6.20. Landslide inventories for areas in central Italy: (a) 16,809 landslides in Umbria–Marche (data set A) identified using aerial photographs taken between 1954 and 1956; (b) 4,233 landslides triggered by rapid snow-melt in January 1997 in Umbria (data set B); and (c) non-cumulative frequency–area distributions of these data showing a common slope of –2.5, suggesting characteristics of self-organized criticality (modified from Guzzetti et al. 2002).

under present conditions. Acceleration of soil erosion is often related to human activity, in particular the clearance of vegetation for agriculture. Van Andel and Runnels (1987) discuss a number of events of this nature,

and Wainwright (1994) considers accelerated erosion to be a major factor in population decline in southwest France in the Bronze Age. The development and consequences of this sort of process are considered in

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Fig. 6.21. Large rock slides in the valley of the River Guadalfeo, Granada, Spain, showing the relationship between slope instability and channel undercutting (photo: John Wainwright).

detail by Wainwright and Thornes (2003). As well as on-site impacts in terms of the impoverishment of soil resources and thus long-term land degradation, there are important off-site consequences, such as pollution of watercourses, flood damage, and reservoir siltation (Meddi and Morsli 2001; De Vente and Poesen 2005) (see Chapters 8 and 20).

Guzzetti (2000) has estimated that from 1410 to 1999, at least 840 landslide events were responsible for 10,555 deaths in Italy (Figure 6.23). The sequence is made up of numerous small events punctuated by a number of major events. For example, 400 people died as a result of the failure of a landslide dam and debris flow in the Passer valley in 1419; 1,200 perished as a

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(a)

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First time failure Earthflow reactivation

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result of a rockslide at Piuro in 1618; 1,917 were killed by the Vaiont event of 1963 (although some sources say the figure was closer to 2,500). The latter event was the endpoint of a complex series of events following the construction of a dam in the Piave valley, Veneto (Petley 1996). The site is a deep gorge, where detailed mapping before the event revealed a number of older slides, although most of these were not considered to be in danger of reactivation (Semeza and Ghirotti 2000). As the dam construction progressed and reservoir levels started to rise to 590 m a.s.l., there was a minor landslip in March 1960 (Figure 6.24), followed by the onset of creep of slope material in the following months as the reservoir gradually filled. Major fractures occurred within the slope material as a result of loading from the water, leading to a major failure producing 700,000 m3 of sediment in November 1960, with waves of up to 20 m in height being produced against the dam. As a result, the water level in the reservoir was carefully lowered from

Fig. 6.22. Idealized evolution of (a) ground-surface properties and (b) surface instability (factor of safety) based on analysis of earthflows in the Barcelonette basin, France (modified from Maquaire et al. 2003), showing the likelihood of periodic reactivation after failure for a period of up to about 150 years.

the failure height of 650 m to a new level of 600 m a.s.l. in January 1961. It was hoped that creep would produce stabilization of the slope material, and a bypass tunnel was also constructed to help drain the slopes. Following this construction in October 1961, reservoir levels were increased again, reaching 700 m a.s.l. in December 1962. Slope movement increased rapidly to 15 mm day−1 at this stage, so water levels were again reduced to 650 m, which again stabilized movement. It was believed that slope materials had become stable, so water levels were increased once again in April 1963, and slow movements started to occur only at water levels of around 700 m, which appeared to support the idea. Rapid movements started to occur at a height of 710 m a.s.l., but the decision to reduce water levels was delayed, and despite commencing in September 1963, there was a catastrophic slip on 9 October 1963, which displaced 250 M m3 of sediment at velocities up to 30 m s−1 . About 30 M m3 of water was displaced by this sediment,

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(a) Vajont Oct.9 1963

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Piuro Sept. 4 1618

Stava July 19 1985

1980

overtopping the reservoir with waves reaching up to 950 m a.s.l., with consequent flooding of downstream settlements leading to the fatalities noted above. Semeza and Ghirotti (2000) emphasize the need for detailed mapping of local conditions, including the presence of older failures, and indeed landslide-susceptibility mapping is an increasingly common planning tool in the minimization of landslide hazards (e.g. Carrara et al. 1991; Gokceoglu et al. 2005). Detailed monitoring using GPS and ground-based techniques is also becoming important, as shown in the assessment of hazards in the Super Sauze earthflows in the French Alps (Maquaire et al. 2003).

The Wider Context In Chapter 1, the link between uplift and erosion has already been touched upon. It is important to recognize

0 2000

Fig. 6.23. Analysis of landslide events that resulted in fatalities in Italy comparing (a) the period 1410–1999 and (b) 1900–99 (modified from Guzzetti 2000).

that there is a complex interplay between uplift, topography, and erosion rates (Burbank and Anderson 2001). Lewis et al. (2000) have suggested that the continued uplift (despite crustal extension) of the Catalan coastal ranges is as a result of flexural isostasy caused by erosion of the rifting margin. Increased topography is likely to accelerate erosion rates in this case, producing a positive feedback. Gargani (2004) has argued that the topography of the Rhône valley in southern France is a result of complex isostatic compensation during the Messinian Salinity Crisis due to unloading of the seabed as sea levels dropped, and due to the consequent incision of the deep river gorge. At a larger scale, Kuhlemann et al. (2002) estimate that erosion of the Alps following uplift has produced 925,800 km3 of sediment over the last 35 Ma. Isostatic compensation for removal of material causes continued uplift, and these figures suggest that almost 7 km of uplift may be the result of the isostatic compensation from this erosion. Similarly, the Rhône valley

(a)

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m 2000 1500 1000 500

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Fig. 6.24. Details of the Vaiont landslide disaster: (a) location map; (b) geological cross-sections before and after the major event of 9 October 1963; and (c) comparisons of precipitation, reservoir level, rate of movement, and subsurface water levels during the construction, leading up to the failure event (modified from Semenza and Ghirotti 2000).

Weathering, Soils, and Slope Processes

will have subsided in compensation for the 30,000 km3 deposited over this time period, while deposition in the Gulf of Lions is of the order of 267,000 km3 and in the Adriatic, of the order of 104,700 km3 , with subsidence in both cases having important feedbacks to coastal and fluvial processes. It has also been argued that the relative rates of plate convergence compared to erosion is an important factor in controlling the shape of orogens (Willett et al. 2001), while Whipple and Meade (2006) have argued that erosion rates can even control plate movements by means of these feedbacks. Weathering and soil erosion can also have major impacts on the carbon cycle. Geologically, these processes can feed back to large-scale climate changes (as discussed for the Himalaya, Tibet, and Andes by Raymo et al. 1988, who argued that uplift would increase weathering rates, leading to a drawdown of atmospheric CO2 and thus to global-scale cooling). Lal (2005) has argued that the state of knowledge is presently incomplete as to whether soil erosion acts as a sink or a source for carbon, with important consequences for ongoing climate change. However, estimates that burial of eroded sediments on land can explain the missing carbon from the global cycle (Stallard 1998) is based on an erroneous assumption about the extent of such burial, related to the issues of scaling erosion rates already mentioned above (Parsons et al. 2004). Alternative hypotheses suggest that erosion ultimately provides a net source of carbon to the atmosphere, although Lal (2003) makes the same erroneous assumption as Stallard in the amount of total erosion that can take place. Whichever of these theories proves to be correct, the Mediterranean as a dynamically eroding area is likely to be a major area of concern for global carbon cycles—for example, Roose (2004) summarized data for carbon erosion from sites in Algeria and other semi-arid areas, and suggested that yields of carbon could be significant, especially on agricultural soils and on sparsely vegetated, steep slopes–and is potentially a key area for attempting to mitigate climate-change impacts in a holistic way.

Conclusion Several decades of research employing a processbased perspective have produced important advances in the understanding of the operation of hillslopes in the Mediterranean region. It is important to recognize the variability inherent within these systems and that more detailed understanding allows the potential for appropriate management of these environments. This vari-

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ability implies that perspectives that are characterized by ‘typical Mediterranean’ conditions are likely to lead to problems, and that a bottom-up (local) approach is required to understanding patterns of slope processes, be they weathering, water erosion, or mass movements. The reflection of different timescales in processes highlights the vulnerability of these systems, especially the relatively slow rate of soil production compared to erosion rates, notwithstanding the problems highlighted with the measurement of the latter. What is now required is a more holistic perspective, in which information about these different processes is integrated—including over longer timescales—and issues of compatibility addressed. This chapter has highlighted a number of (but by no means all) important non-linearities and scaling issues within and between different processes. The evaluation of these non-linearities and their impacts must remain a priority in the investigation of Mediterranean landforms. They are significant in understanding the evolution of the Mediterranean, and will become increasingly important in the sustainable management of these sensitive environments, with regional and global implications.

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van Wesemael, B., Verstraten, J. M., and Sevink, J. (1995), Pedogenesis by clay dissolution on acid, low-grade metamorphic rocks under Mediterranean forests in southern Tuscany (Italy), Catena 24: 105–25. Vavliakis, E. G., Haristos, D. A., and Balafoutis, C. (1990), Indirect influence of man-made factors on the dissolution rate of dolomitic marble in Thessaloniki area (northern Greece), Zeitschrift für Geomorphologie 34: 475–80. Wainwright, J. (1994), Anthropogenic factors in the degradation of semi-arid regions: a prehistoric case study in southern France, in A. C. Millington and K. Pye (eds.), Effects of Environmental Change on Drylands. John Wiley & Sons, Chichester, 285–304. (1996a ), Constitution et évolution des nappes archéologiques, in J. Gascó, L. Carozza, S. Fry, R. Fry, J.-D. Vigne, and J. Wainwright (eds.), Le Laouret et la Montagne d’Alaric à la Fin de l’Âge du Bronze, un hameau abandonné entre Floure et Monze (Aude). Centre d’Anthropologie, Toulouse, 251–60. (1996b), Infiltration, runoff and erosion characteristics of agricultural land in extreme storm events, SE France, Catena 26: 27–47. (1996c), Hillslope response to extreme storm events: the example of the Vaison-la-Romaine event, in M. G. Anderson and S. M. Brooks (eds.), Advances in Hillslope Processes. John Wiley & Sons, Chichester, ii. 997–1026. (1996d ), A comparison of the infiltration, runoff and erosion characteristics of two contrasting ‘badland’ areas in S. France, Zeitschrift für Geomorphologie Suppl. 106: 183–98. and Thornes, J. B. (2003), Environmental Issues in the Mediterranean: Processes and Perspectives from the Past and Present. Routledge, London. Mulligan, M., and Thornes, J. B. (1999), Plants and water in drylands, in A. J. Baird and R. L. Wilby (eds.), Ecohydrology. Routledge, London, 78–126. Parsons, A. J., Müller, E. N., Brazier, R. E., Powell, D. M., and Fenti, B. (2008), A transport-distance approach to scaling

erosion rates: 1. Background and model development. Earth Surface Processes and Landforms 33: 813–26. Whipple, K. X. and Meade, B. J. (2006), Orogen response to changes in climatic and tectonic forcing, Earth and Planetary Science Letters 243: 218–28. Willett, S. D., Slingerland, R., and Hovius, N. (2001), Uplift, shortening, and steady state topography in active mountain belts, American Journal of Science 301: 455–85. Wise, S. M., Thornes, J. B., and Gilman, A. (1982), How old are the badlands? A case study from southeast Spain, in R. B. Bryan and A. Yair (eds.), Badland Geomorphology and Piping. GeoBooks, Norwich, 259–77. Wondafrash, T. T., Sancho, I. M., Miguel, V. G., and Serrano, R. E. (2005), Relationship between soil color and temperature in the surface horizon of Mediterranean soils: a laboratory study, Soil Science 170: 495–503. Woodward, J. C. and Goldberg, P. (2001), The sedimentary records in Mediterranean rockshelters and caves: archives of environmental change, Geoarchaeology 16: 327–54. Macklin, M. G., and Lewin, J. (1994), Pedogenic weathering and relative-age dating of Quaternary alluvial sediments in the Pindus Mountains of northwest Greece, in D. A. Robinson and R. B. G. Williams (eds.), Rock Weathering and Landform Evolution. John Wiley & Sons, Chichester, 259–83. Yaalon, D. H. (1997), Soils in the Mediterranean region: what makes them different? Catena 28: 157–69. Yair, A. (1995), Short and long term effects of bioturbation on soil erosion, water resources and soil development in an arid environment, Geomorphology 13: 87–99. Yassoglou, N., Kosmas, C., and Moustakas, M. (1997), The red soils, their origin, properties, use and management in Greece, Catena 28: 261–78. Zalidis, G., Stamatiadis, S., Takavakoglou, V., Eskridge, K., and Misopolinos, N. (2002), Impacts of agricultural practices on soil and water quality in the Mediterranean region and proposed assessment methodology, Agriculture Ecosystems and Environment 88: 137–46.

This chapter should be cited as follows Wainwright, J. (2009), Weathering, soils and slope processes, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 169–202.

7

Vegetation and Ecosystem Dynamics Harriet Allen

Introduction Within the Mediterranean region a number of distinctive vegetation communities can be recognized, comprising some 25,000 species, of which about 50 per cent are endemic. Broadly defined, these originated with the establishment of a mediterranean-type climate about 3.2 million years ago, since when they have been subject to the vicissitudes of glacial–interglacial climate changes, plus the intensification of human impact during the last 10,000 years (Chapters 4 and 9). These communities are dynamic, responding to environmental changes at a variety of scales, both spatial and temporal. This chapter explores the characteristics of these communities and examines the relationships between ecosystem dynamics and biodiversity, and ecosystem response to disturbance. For example, each year fires burn out of control and are the subject of regular news stories during summer months. While fires may be economically devastating and lead to loss of life (Chapter 19), ecologically their incidence is an important dynamic component of Mediterranean ecosystems and may, indeed, be crucial to the successful propagation and spread of plants and communities regarded as typically Mediterranean. Associated animal populations generally recover quickly despite inevitable loss of life in some populations. Thus understanding the role of fire and other disturbance factors such as grazing is key to understanding Mediterranean vegetation communities and ecosystem dynamics. The chapter concludes with an evaluation of the likely response of vegetation communities to potential atmospheric and land use changes.

Vegetation Communities While a number of distinct vegetation communities have been identified, a common characteristic is an ability

to survive hot, dry summers and cool, wet winters, together with frequent disturbances. Many of the communities are dominated by shrubs, and Mediterranean evergreen sclerophyllous shrublands are recognized as one of the defined ecosystems of the world (di Castri 1981). Such shrublands are at the centre of a continuum of communities which vary along gradients of moisture availability, temperature, and nutrient availability, usually determined by substrate, and human activity (Figure 7.1). At the extreme ends of these gradients, but still Mediterranean, are sclerophyllous woodlands, coniferous and deciduous forests, savannas and grasslands grading into steppe and semi-desert shrublands, and heathlands. The relationship between vegetation and climate is recognized in bioclimatic classifications (Quézel 1985). Eight different life zones (Figure 7.2) are identified according to temperature and precipitation indices, varying with either latitude or altitude, and therefore recognizable in a north–south traverse across the Mediterranean basin or along an altitudinal transect. In each of these bioclimatic zones, some of the dominant species can be identified as characteristic bioindicators.

Evergreen Sclerophyllous Shrublands Many Mediterranean countries have their own names for evergreen shrubland communities, with the terms ‘maquis’ or ‘matorral’ widely used (Table 7.1). This can cause confusion in the literature, as the names are not defined precisely (Allen 2001), and may be compounded by classifications that depend not just on the life form of the species present (e.g. whether tree or shrub, or perennial or annual), but also on the height and proportion of ground cover. For example, maquis communities are generally regarded as comprising relatively tall-growing shrubs (Figure 7.3), while garrigue is a

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Harriet Allen Savannas and grasslands Desert and semi-desert shrublands

Fertilization

Greater nutrient availability

Aridity Climatic/ environmental control Human intervention Desertification

Evergreen sclerophyllous shrublands

Coniferous and deciduous forests

Higher rainfall, lower winter temperature

Greater moisture availability

Low nutrient availability

Sclerophyllous woodlands

Heathlands

Fig. 7.1. Mediterranean vegetation communities and trajectories of change.

lower-growing community (Figure 7.4), but a continuum exists between the two. Mediterranean shrublands are dominated by sclerophyllous taxa. These have tough, leathery evergreen leaves, with a low surface-area to volume ratio, thick cuticles, and stomata sunken into grooves on the underside of the leaf. These characteristics are widely believed to be adaptations to water stress, as a means to control transpiration losses. Other selective forces might include response to fire, herbivory, and nutrient-poor substrates (Mooney and Dunn 1970). In addition to evergreens, deciduous shrubs and aromatic plants are common in shrubland communities. Some deciduous shrubs show seasonal dimorphism; that is, leaf morphology occurs in two forms as a response to intense summer drought. Rates of photosynthesis and respiration are higher in winter leaves due to higher concentrations of soluble sugars, chlorophyll, fat, nitrogen, and free and bound amino acids compared with summer leaves (Margaris 1977). Within a genus some species may be evergreen, while others are deciduous. For example, there are about ten deciduous Pistacia species, but only one evergreen, Pistacia lentiscus. Among the Mediterranean oaks, there are six evergreen species, including Quercus ilex, Q. suber, and Q. coccifera, about 35–40 deciduous, including Q. robur, Q. cerris, and Q. pyrenaica, and one semi-deciduous, Q. infectoria (Blondel and Aronson 1999). The Mediterranean flora contain some 49 per cent of the world’s aromatic taxa, which produce essential

oils. The majority belong to the mint (Lamiaceae), carrot (Apiaceae), and daisy (Asteraceae) families. The production of essential oils within leaves appears to peak in summer (Ross and Sombrero 1991), which suggests a regulatory factor associated with seasonality. This need not be drought, but might be a defence mechanism against herbivory, either by insects or grazing animals. The leaves of many aromatic plants are lacking in nitrogen, and as soils are often nitrogen- and phosphoruslimiting, leaf growth is slow. Production of long-lived leaves is therefore an advantage. As well as being a deterrent to herbivory, essential oils may be anti-microbial or anti-fungal, may attract pollinators by mimicking insect pheromones, and may cool leaves through volatilization of the oils, thus reducing transpiration. Maquis communities are the most widespread of the mediterranean ecosystems. They form a narrow coastal belt and penetrate inland according to climatic and topographic variability, being more extensive in the western than eastern Mediterranean (Figure 7.5). Their altitudinal and latitudinal limits appear to coincide with an average minimum temperature in the coolest month of 0◦ C. Summer maximum temperatures do not appear to be a limiting factor. Their extent usually coincides with mean annual precipitation levels of 350 to 1,500 mm, though in North Africa and the Middle East the limiting annual total is less than 200 mm. While nutrient availability may not be directly responsible for the distribution of maquis communities, they are absent from salty soils, and also from thin, highly eroded soils

Vegetation zonation

Vegetation communities

Cryo-Mediterranean

Alpine on rock, scree, and gravel

Alti-Mediterranean

Sub-alpine grasslands, herbaceous perennials, and dwar f junipers

Juniperus , Bromus, Festuca, Poa, Phleum

Oro-Mediterranean

Coniferous woodland

Pinus uncinata, Pinus mugo

Montane Mediterranean

Deciduous woodland

Fagus sylvatica; conifers (Pinus, Cedrus, Abies), Juniperus

Supra-Mediterranean

Meso-Mediterranean

Infra-Mediterranean (Western Morocco)

4,000

Saxifraga, Androsace, Aubretia Average minimum temperature of the coldest month (°C)

3,000

– 9°C

Quercus humilis, Quercus cerris, Deciduous oak forests Quercus frainetto, Quercus macedonia, and semi-deciduous in Ostrya carpinus, Carpinus orientalis, Nor th Africa and Spain Corylus sp ,Tilia sp., Fraxinus ornus, Acer sp.s Quercus ilex, Quercus calliprinos, Pinus halepensis, Pinus brutia. Where heavy Evergreen oak woodands anthropogenic activity: Quercus coccifera, and shrublands Calycotome villosa, Genista acanthoclada Olea europea, Ceratonia siliqua, Phillyrea Dense coastal woodand; media, Pistacia lentiscus, Laurus nobilis, sclerophyllous, evergreen; Quercus suber, Pinus pinaster, significant human impact Chamaerops humilis

2,000

–5°C –2°C +1°C

1,000 +3°C

Argania spinosa and Acacia gummifera

+5°C

+7°C

0 30

32

34

36 38 40 Latitude (°N)

Fig. 7.2. Bioclimatic life zones of the Mediterranean region (after Blondel and Aronson (1999) and Le Houérou (1990)) (see Chapters 5 and 23).

42

44

46

Vegetation and Ecosystem Dynamics

Thermo-Mediterranean

Altitude (metres)

Bioindicator taxa

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TABLE 7.1. Terms used to describe Mediterranean sclerophyllous shrubland High matorral Israel Greece Italy Macchia alta France Maquis

Spain

Matorral denso, espinal

Middle matorral

Macchia bassa Garrigue (sometimes only with reference to calcareous soils) Matorral claro, jaral (on siliceous soils)

Low matorral Batha Phrygana Garriga Tomilar

Garrigue (or Landes on siliceous soils)

Source: After Tomaselli (1981a) and Margaris (1981).

at the climatic limits of their growth; a likely response to low soil moisture retention. Thus at higher elevations, they are found in association with calcareous soils rather than on poorly developed metamorphic-derived soils (Quézel 1981a, b). Typical evergreen or semi-evergreen trees and shrubs that have a circum-Mediterranean distribution (Figure 7.6) include olive (Olea europea), carob (Ceratonia siliqua), Pistacia, and some members of the Cistaceae family (Figure 7.7), e.g. Cistus salvifolius, and Ericaceae family, such as Erica arborea and Arbutus unedo. Many aromatic labiate species are also widespread, such as rosemary (Rosmarinus officinalis), lavender (Lavandula spp.), and thyme (Thymus spp.). Some taxa have more restricted distributions. Those more common in the western Mediterranean include species of Genista, Cytisus, and Ulex (all members of the pea family— Fabaceae), and several members of the Cistaceae and Ericaceae families. In the eastern Mediterranean, there are several labiate species of Phlomis, Satureja, and Salvia, and a greater number of species of shrubs such as Rhamnus and Daphne than are found in the western Mediterranean. The Judas tree (Cercis siliquastrum), though widely planted for ornament elsewhere, has its origin in the eastern Mediterranean. Four species of oak and two pines are widely found in sclerophyllous shrublands: holm oak (Quercus ilex), kermes oak (Q. coccifera), cork oak (Q. suber), and Aleppo pine (Pinus halepensis) in the western Mediterranean, and Q. calliprinos and Calabrian pine (P. brutia) in the eastern Mediterranean (Figure 7.6). These are examples of vicariant or disjunct distributions; that is, the distributions of two closely related but non-coexisting species. Conifers that grow in association with sclerophyllous shrublands include junipers (Juniperus phoenicia, J. oxycedrus). There is some indication that arborescent taxa are more common in the eastern than western Mediterranean (Quézel 1981a).

Sclerophyllous Woodlands Although evergreen oaks such as holm oak (Quercus ilex) may be present in the sclerophyllous shrubland communities, at the wetter and cooler margins of the Mediterranean they can form more or less continuous woodland cover (Terradas 1999). Q. ilex has a wide distribution, but is replaced in the eastern Mediterranean by Q. calliprinos. Holm oak forests are particularly extensive in Sicily, Italy, Corsica, France, Spain, and Morocco and, historically, have been of considerable economic importance, in particular as a source of charcoal (Terradas 1999). Pollen diagrams from Huelva in south-western Spain suggest that woodland management occurred as early as 6,000 cal. years BP (Harrison 1996), although in other areas of the Mediterranean basin such management might date from later periods.

Savanna Woodlands In addition to closed woodland, savanna-type landscapes of ‘trees without forests’ can be recognized across the Mediterranean—areas of trees scattered amongst other vegetation types. Grove and Rackham (2001) regard their origin on the Iberian Peninsula to be the result of management formalized over centuries to create dehesas (or montados in Portugal). Such a cultural origin is likely elsewhere but savanna woodlands also occur in the drier areas of the Mediterranean. Trees of savanna systems were traditionally managed by coppicing and pollarding and their continued existence was economically crucial: dehesas traditionally provided acorns and shelter for livestock, tannin, cork, and firewood, and pasture for sheep and cattle (Joffre 1992). Trees associated with Iberian dehesas include the evergreen oaks (Quercus ilex, Q. rotundifolia) and the cork oak (Q. suber) and deciduous oaks (Q. faginea and Q. pyrenaica). Conifers and deciduous trees are also characteristic of savanna woodlands. In the Alpujarra (Spain), Italy, Corsica, and western Crete, sweet chestnut (Castanea sativa) is important, as are beech (Fagus sylvatica) in Italy, Corsica, and northern Greece, and carob (Ceratonia siliqua) in southern Portugal and Crete (Grove and Rackham 2001). The emergence and longevity of savanna-type woodlands has resulted in high levels of biodiversity and they are often important conservation sites. But their continued existence, and that of associated fauna, is threatened by land use changes (Chapter 23).

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Fig. 7.3. Invasion of sclerophyllous maquis vegetation into an old olive orchard, Crete (photo: Harriet Allen).

Coniferous Forests In mountainous areas and on coastal sandy soils, coniferous forests are widespread, although plantations also exist at lower elevations. The most widespread conifers are the pines, but other important taxa include cedars (Cedrus atlantica in the Atlas Mountains, C. brevifolia in Cyprus, and C. libani in Lebanon and Turkey) and, in high mountain regions, firs (Abies sp.). About ten different pine species grow around the Mediterranean, although only four are regarded as true Mediterranean species (Klaus 1989): Aleppo pine (Pinus halepensis), Calabrian pine (P. brutia), Canary Island pine

(P. canariensis), and stone or umbrella pine (P. pinea). Of these, Aleppo and Calabrian pines are another example of vicariant distributions (Figure 7.6). Genetically and ecologically similar, they coexist in two small areas of Greece, in south-eastern Anatolia, and in Lebanon and Israel, where they form natural hybrids (Barbero et al. 1998). Geological, archaeological, and historical evidence points to an eastern origin for both and climate is the likely determinant of their modern distribution (Biger and Liphschitz 1991). P. halepensis requires a more humid and warmer climate; it cannot survive temperatures of −10◦ C. It therefore spread west in response

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Harriet Allen

Fig. 7.4. Typical garrigue vegetation in Crete. Note the lower-growing, rounded nature of the grazed shrubs (photo: Harriet Allen).

to its climatic needs and climatic change. By contrast, P. brutia can survive lower temperatures and therefore remained in its area of origin. However, its occurrence in Israel and the surrounding areas may be a recent phenomenon, given its absence from wood samples collected at Israeli archaeological sites during the late Holocene (Biger and Liphschitz 1991). Pine forests are estimated to cover about 5 per cent of the total Mediterranean, but comprise about 25 per cent of the forest cover in Anatolia and up to 75 per cent in North Africa (Barbero et al. 1998). Their extent has changed considerably in recent decades. In North Africa,

decline has followed clearance for cultivation, felling of timber for construction and charcoal production, overgrazing, and frequent fires. By contrast, in the European Mediterranean, pine forest has increased in area as a result of land use changes, which include afforestation, and the ability of pine to invade abandoned or burned land (Barbero et al. 1998).

Grasslands, Steppe, and Semi-desert Shrublands Around the Mediterranean basin, grasslands tend to occur where mean annual precipitation is less than

Vegetation and Ecosystem Dynamics

0

209

500 km

Fig. 7.5. Extent of maquis communities across the Mediterranean region (after Quézel 1981a).

about 300 mm (Blondel and Aronson 1999). In addition, perennial grasses form a significant component of ‘alti-mediterranean’ or subalpine communities (Figure 7.2). Some of the most extensive grasslands are found in North Africa and Spain, such as those dominated by the perennial bunch grass, Stipa tenacissima (alfa grass). This is a slow-growing drought- and high-temperature endurer, which mainly regenerates vegetatively and grows as tussocks. Between the tussocks, annual grasses invade. These persist until the Stipa tussocks have encroached into the inter-tussock spaces, while the older tussocks become senescent. The dying tussocks then form new inter-tussock areas, to be invaded by annual grasses. This two-dimensional regeneration pattern is believed to confer stability on the Stipadominated grasslands; the presence of annuals, albeit at different densities and biomass, ensures that few areas of bare ground exist for long, which reduces the potential for soil erosion (Clark et al. 1998; Chapter 6). In especially arid areas, such as parts of Spain and North Africa, Stipa and Brachypodium grasses grow in association with shrubs such as wormwood (Artemisia). In more arid areas, grassland species grow in association with other arid taxa, such as Artemisia, Pistacia, and Ephedra. These are identified as being Irano-Turanian

in origin; that is, their centres of origin are the semiarid steppes of central Asia, where summers are exceptionally hot and winters very cold and dry. Today, the Irano-Turanian phytogeographical region extends as far west as Jordan and the non-coastal parts of Lebanon, Israel, and Syria. Formerly, during the glacial stages of the Quaternary it spread further west across the Mediterranean region. Thus the steppe communities, identified by pollen analysis, at the height of the global Last Glacial Maximum in northern Europe, about 18,000– 20,000 years ago, were indicative of extensive Mediterranean aridity. A large number of Irano-Turanian elements are recognized within the contemporary Mediterranean flora (Blondel and Aronson 1999) and these are believed to represent the successful ‘invasions’ of xeric plants into more humid, mediterraneantype communities following human activity; disturbed ground is susceptible to invasion by weedy taxa more representative of the drier habitats of the Mediterranean. This is referred to by Blondel and Aronson as Zohary’s ‘law’ based on the ‘expansion drive’ of desert plants into the Mediterranean region of the Near East (Zohary 1962). This expansion of grasses and associated taxa underlies the recognition of Mediterranean grasslands and steppe as anthropogenic (Figure 7.8), being

210

Harriet Allen Olea europea subsp. oleaster

0

Arbutus unedo

1000 km

Cistus salvifolius

Lavandula stoechas

Quercus ilex Quercus calliprinos

Pinus halepensis Pinus brutia

Quercus suber

Cercis siliquastrum

Fig. 7.6. Distribution maps for circum-Mediterranean taxa, Olea europea subsp. oleaster (wild olive), Arbutus unedo (strawberry tree), Cistus salvifolius (sage-leaved cistus), Lavandula stoechas (lavender), and vicariant taxa, Quercus ilex and Q. calliprinos (holm oaks), Pinus halepensis (Aleppo pine) and P. brutia (Calabrian pine), Quercus suber (cork oak), and Cercis siliquastrum (Judas tree). Based on Daget (1980), Quézel (1985), and Biger and Liphschitz (1991).

Vegetation and Ecosystem Dynamics

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Fig. 7.7. Cistus ladanifer-dominated scrub vegetation of the Algarve, Portugal (photo: Harriet Allen).

degraded maquis communities. In areas of hyper-aridity, where saline soils form due to high water tables, salttolerant taxa grow, many of which are members of the Chenopodiaceae (goosefoot) family.

Heathlands Heathlands may occur where nutrient availability is low, for example on poor siliceous soils. Although Mediterranean heathlands received only limited mention in the review of global heathland ecology (Specht 1979), they are recognized in phytosociological surveys by their typical vegetation associations, especially on the sandy soils of France, known as landes. Typical genera are Erica, Cistus, Quercus, and Genista. Sandy soils are generally acidic, which limits the availability of phosphorus, compared with neutral to alkaline soils. When such nutrient-poor soils are fertilized,

productivity increases (di Castri 1981) which demonstrates the nutrient-limiting characteristics of Mediterranean heathlands. The extent to which Mediterranean heathlands are natural or a result of human impact is debatable. Pollen analysis indicates that they can occur spontaneously or have a long history. Low-growing shrub formations of Pistacia, Phillyrea, Juniperus, and Thymelaea have been identified in communities that lacked trees such as oak (Quercus), fir (Abies), or beech (Fagus). This suggests the heath taxa were not merely representative of woodland understorey communities, but existed as communities in their own right (Blondel and Aronson 1999). Mediterranean heathlands generally occur as isolated islands surrounded by more calcareous substrates. Floristic surveys of the heathlands of the Strait of Gibraltar region of Spain show them to have high levels of species endemism, though low species richness, due to

212

Harriet Allen Primary maquis/ sclerophyllous evergreen shrubs

Regeneration Reasonable exploitation

Secondary maquis

Degeneration

Excessive exploitation

Garrigue

Grazing

Fire

Cultivation

Steppe

Abandonment

Pasture/grassland Cultivation

Degeneration Regeneration

Regeneration

Abandonment

Orchards/crops

Fig. 7.8. Theoretical degradation and regeneration sequences for primary maquis or sclerophyllous evergreen shrub communities.

the nutrient-poor status of the acid, sandy soils (Ojeda et al. 1995). This could make them special targets for conservation.

Sub-alpine, Alpine, and Cliff Communities As a highly mountainous region, sub-alpine and alpine communities form an important component of Mediterranean flora. In the sub-alpine zone, a common characteristic is that of ‘hedgehog’ plants, a term used to describe their spiny, dwarf form (Figure 7.9). Many of these are members of the Fabaceae (pea) family, such as Astragalus angustifolius (a milk vetch) and Genista acanthoclada, both found in Crete. The ‘hedgehog’ zone is well developed in the Pyrenees, Sierra Nevada mountains of Spain, in the Moroccan Atlas mountains, and in Greece, Crete, and the Taurus mountains of Turkey. Growing alongside these plants are often rich swards of perennial bunch grasses. Where rainfall is heavier, mountain grasslands form important pasture land for summer grazing; transhumance remains an important part of livestock management in the Iberian Peninsula (Ruíz and Ruíz 1986) and in Greece and Crete (Rackham and Moody 1996). Within the sub-alpine zone there may also be shrub forms of trees such as Juniperus communis subsp. nana. Many of these communities have been influenced by fire and grazing.

True alpine communities are found at the highest elevations, for example above 2,600 m in the Pyrenees and Sierra Nevada (Polunin and Smythies 1973) and above 2,200 m on Crete (Sfikas 2002). The typical growth habit of alpine plants is that of prostrate, creeping, rosette, or cushion form sheltering amongst areas of bare rock, boulders, gravel, and scree. Typical alpine genera include Saxifraga, Androsace, and Aubretia. In addition, geophytes (bulbs) are a relatively common element of alpine and subalpine communities. A common characteristic of Mediterranean alpine communities is the high numbers of endemic species—a result of the geographical and reproductive isolation of these high elevation areas—which makes them of great interest to botanists. The Pyrenees have over 180 endemic alpine species (Polunin and Smythies 1973) and a similar number are found on Crete (Rackham and Moody 1996). Cutting through many mountainous areas are deep ravines and gorges, for example the Samaria and Imbros gorges of Crete, and the Vikos Gorge in the Pindus Mountains of Greece. Steep cliffs, with steps and crevices, create strong microclimatic gradients, which, together with factors such as aspect, lead to a variety of habitats and assemblages of plants, many of which are out of reach from browsing and grazing animals. These sites are often rich in endemic species and have frequently acted as refugia for taxa which have undergone migrations and population recessions through glacial– interglacial cycles (Chapter 4). For example, Tertiary

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Fig. 7.9. The typical ‘hedgehog’ shape of the alpine, Euphorbia acanthothamnos, growing among rocky scree of the Psilorites Mountains, Crete (photo: Harriet Allen).

relicts such as Viola cazorlensis, Pinguicula vallisneriifolia, and Ptilotrichum reverchonii are found in the gorges of the Sierra de Cazorla, southern Spain (Polunin and Smythies 1973). The inaccessibility of the high cliff faces of the gorges, which may extend from sea-level up to the sub-montane and montane zones, means that they have remained little touched by human activity and so are important conservation sites.

Wetlands Wetland vegetation communities, of international importance for indigenous wildlife and migrating wildfowl, occur around the Mediterranean. As they differ according to duration of flooding, depth of water column, and salinity, there is a continuum between saline coastal marshes and drier, more freshwater marshland. Notable wetland communities of the

Mediterranean region include the marismas (annually flooded marshlands) of the Coto Doñana in Spain, the Camargue of France, Amvrakikós of Greece, Lakes Bardawil and Burullus in Egypt, and Garet el Ichkeul and Sebkhet Kelbia in Tunisia (Chapter 9). The future of some of the larger wetlands is assured because of their importance to conservation, but many of the smaller ones are transient, both in terms of their ecological status as a response to changing sea levels and because of catchment land use changes which alter their hydrological balance.

‘Semi-natural’ Landscapes Agriculture spread across the Mediterranean from the Near East during early to mid-Holocene times. Consequently few vegetation communities can be regarded as untouched and some have developed alongside

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cultivation practices, leading to ‘semi-natural’ landscapes. Examples include olive groves and terraced landscapes often associated with olives and orchard trees. In Crete traditionally managed olive groves have formed part of the island’s landscape ecology for thousands of years (Allen et al. 2006). Pollen evidence confirms cultivation as early as 5,700 (uncalibrated) years BP, before the beginning of the Early Minoan I period and earlier than on either the Greek or Turkish mainlands (Bottema and Sarpaki 2003). Such groves often support a rich ground flora, which some researchers believe resembles that of ‘natural’ Mediterranean ecosystems (Loumou and Giourga 2003). The ground flora is characterized by annuals and geophytes (bulbs). The annuals survive because regular shallow tilling and disturbance of the soil promotes germination of seeds; among the annuals are a number of species regarded as rare weeds (Phitos et al. 1995). Geophytes survive because they are buried beneath the depth of tillage; for example, species of the Serapias orchid. The diversity of flora provides important habitats for a variety of fauna, such as mammals, birds, and insects (Loumou and Giourga 2003). Terraced landscapes (Figure 7.10) developed as an inevitable response to complex topography, and terraced agriculture is now integral to large areas of the Mediterranean region (Grove and Rackham 2001). Some terracing is ancient and based on the construction and maintenance of retaining walls; other terracing is more recent and relies more on the use of bulldozers. A variety of cultivation can be practised and mixed on terraces, including olives, vines, orchard trees, and arable crops. As with olive groves, some of these can develop speciesrich ‘semi-natural’ vegetation communities. However, maintenance of terraces is often labour intensive and much of it is now falling into disrepair, as terraces are abandoned—a consequence of twentieth-century agricultural intensification on more easily cultivable land (Allen et al. 2006). Abandonment has led, in places, to invasion by perennial shrub and bush species, creating maquis communities (Figure 7.3). The consequence of collapsing terraces has also fuelled a debate about their role in soil erosion, as illustrated by an example from the Alpujarra, in Andalucia. Historically, four periods of neglect of terraces have been identified as responsible for increased soil erosion (McNeill 1992). One of these came after the expulsion of the Moors from Spain in the fifteenth century AD. The Moors had established a system of well-tended terraced and irrigated agriculture, which subsequently declined leading to erosion of topsoil and its deposition in the deltas of the Mobril and Adra rivers. More recently, population decline in the region and abandonment of terraces, following

agricultural deintensification since the Second World War, are believed to have led to increased soil erosion (McNeill 1992). However, there is some uncertainty about rates of erosion and sediment yield in the Alpujarra, as these are measured as having a high degree of spatial variability (Douglas et al. 1994). This makes it difficult to extrapolate them and generalize with respect to the role of terracing. Indeed there is a largely positive view that terrace abandonment may reduce runoff and erosion of topsoil as maquis vegetation increases in extent (Thornes 1998). The diversity of Mediterranean biomes and vegetation communities means that there are numerous ecotones, or transitional habitats formed by the overlapping areas of two adjacent communities. These may be small in area, but this adds to the importance of ecotones and serves to increase the spatial heterogeneity of Mediterranean communities—there are large differences in animal and vegetation communities across short distances. The locations of ecotones are determined by factors such as climate, soils, and disturbance regimes, for example grazing and fire. Such variability adds considerably to landscape diversity and is a reflection of the ecosystem dynamics occurring within different communities.

Plant Biodiversity and Ecosystem Dynamics Empirical research, in a variety of world biomes, suggests that there are strong reasons to assume that biodiversity has an influence on ecosystem dynamics, although trying to establish the relationships between the two is difficult, especially as some of these relationships might be multi-variable in cause and scale-dependent. Assessing the possible influence of biodiversity on ecosystem stability helps to inform the debate about the fragility/resilience of Mediterranean ecosystems. The consequences of this need to be assessed as models of global biome change predict that mediterraneantype ecosystems will probably experience a large loss of biodiversity through the twenty-first century (Sala et al. 2000; and Chapter 23). It may be that Mediterranean ecosystems are becoming (even) more vulnerable to change. However, this has to be examined in the context of Mediterranean landscapes and ecosystem stability, which have traditionally been regarded as vulnerable but are now coming to be recognized as relatively robust. Grove and Rackham (2001) illustrate this debate in a discussion of the question of ‘ruined landscapes’: millennia of human activity have supposedly led to degradation of Mediterranean landscapes,

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Fig. 7.10. Olive terraces on Crete (photo: Harriet Allen).

when in fact many plant communities and landscapes have developed under different patterns of human occupancy and variable climate, to the extent that they have become quite resilient. It is likely that this resilience is promoted by the region’s rich floristic diversity. Hence there is a need to examine why the region is so floristically rich (see also Chapter 23) and to assess the research on the influence of biodiversity on ecosystem processes and stability in the Mediterranean context. There are an estimated 25,000 species of flowering plants and ferns in Mediterranean Europe, of which about 50 per cent are endemic, in an area of about 2.3 million km2 , compared with just 6,000 species in the 9 million km2 of non-Mediterranean Europe (Quézel 1985). Nested within the region are nine diversity hotspots identified on the basis of numbers of endemic plants that appear to be threatened. These are the High and Middle Atlas Mountains, The Betic-Rif complex,

the Maritime and Ligurian Alps, the Tyrrhenian Islands, southern and central Greece, Crete, Anatolia and Cyprus, Syria–Lebanon–Israel, and Mediterranean Cyrenaica in Libya (Médail and Quézel 1997). No single explanatory factor for species richness can be advanced, as the environmental variables predominantly responsible apply at different spatial and temporal scales (see for example Willis and Whittaker 2002). For the Mediterranean region relevant factors include geological and palaeogeographical factors (such as the Alpine mountain building episode leading to reproductive isolation of populations), the consequences of environmental changes (glacial–interglacial cycles), the importance of ecological stress and competition between species, the diversity of habitats (especially relating to topography, microtopography, and microclimate), and the role of disturbance (for example, fire).

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Geological and palaeogeographical factors are normally cited to explain continental- to regional-scale diversity patterns. For example, some taxa that are typically regarded as Mediterranean, in fact evolved prior to the establishment of a mediterranean-type climate: pollen grains from olive (Olea), Phillyrea, Cistus, and Pistacia, are present in off-shore cores from the western Mediterranean Sea in deposits older than 3.2 million years (Suc 1984; Chapter 4). Some mediterranean-type taxa also appear to have their origins in non-Mediterranean regions, for example asparagus (Asparagus), carob (Ceratonia), and oleander (Nerium), which are believed to have spread from Africa at a date earlier than the Oligocene and Miocene (Raven 1973). The location of the Mediterranean region, at the crossroads of Europe, Asia, and Africa, has facilitated the spread of flora from other regions. Other historical explanations for an enriched regional biota are based on the response of taxa to environmental changes, such as those associated with the glacial–interglacial cycles of the Quaternary. The forests of southern Europe acted as refugia for tree species from northern Europe during glacial periods. Around Lake Ioannina in north-western Greece, for example, deciduous forest communities included oak (Quercus), elm (Ulmus), hornbeam (Carpinus betulus), and oriental hornbeam (C. orientalis) which had disappeared from the forests of northern Europe (Tzedakis 1993). Greater diversity in the Mediterranean applies at the genetic as well as at the specific level. Genetic differentiation of tree populations shows higher maternally inherited chloroplast DNA among the refuge populations of southern Europe (Thompson 1999; Hewitt 2000). However, greater diversity is not necessarily found in all measures of genetic diversity in refugial populations (Comps et al. 2001), which complicates the interpretation of molecular evidence for refugia. The varied topography and tectonic history of the Mediterranean region, which have created isolated populations, also help to explain plant species richness and patterns of endemism at regional to subregional and local scales, particularly in mountainous areas. Geographical isolation can lead to reproductive isolation allowing allopatric speciation and the evolution of endemics to occur. On Crete 26.6 per cent of the plant species found at elevations higher than 1,500 m in the Lefka Ori, or White Mountains, are endemic to the island; nested within this, 10 per cent are endemic to individual mountain massifs within the range (Vogiatzakis et al. 2003). The considerable topographic variability, especially in mountainous areas,

means that there are steep ecological gradients, providing a diversity of physical habitats for plants and animals. This promotes species richness at the local to landscape scales. Nested within this are local scale, short-term processes, resulting from factors such as fire and grazing, which are discussed in more detail below. From this brief review of reasons for biodiversity patterns it should be apparent that variables accounting for species richness differ with scale, both temporal and spatial, and that this needs to be taken into consideration when trying to examine the potential role of biodiversity in ecosystem dynamics.

Diversity and Ecosystem Functioning Identifying the components of biodiversity that are related to ecosystem functioning is currently the subject of much debate (see e.g. Schulze and Mooney 1994). In part this is because of the difficulty of defining the term ‘biodiversity’ and in part because of the difficulty in establishing what to measure. One of the commonest usages of the term is to mean ‘species richness’, and it is generally believed that greater plant species richness leads to greater productivity in plant communities, which leads to greater nutrient retention, and greater ecosystem stability (see e.g. Tilman 2000). As yet there have been few empirical studies across a range of ecosystems to test the relationships between diversity and stability. However, some results are being derived from grassland manipulation experiments in Europe, including the Mediterranean region. These suggest that species richness contributes to net primary productivity, but that so too does the type of species grown, i.e. whether perennial or annual, or whether the grassland herbaceous seed mix includes nitrogen-fixing taxa. The results of grassland manipulation experiments revealed a positive relationship between species richness and above-ground primary productivity (biomass) when measured in eight field trials across Europe. A reduction in the number of species grown in a plot resulted in a log-linear reduction in biomass (Hector et al. 1999), although the results were less conclusive for the Mediterranean field plots in Portugal and Greece compared with those in northern Europe. Experiments manipulating the type of grass grown at the sites on Lesvos, Greece, attempted to assess the role of functional diversity through inclusion or exclusion of perennial and annual grasses (Troumbis et al. 2000). Plots containing the perennial grass, Phalaris coerulescens, recorded the greatest primary productivity, suggesting that it may

Vegetation and Ecosystem Dynamics

be the presence of particular types of species that could be a determinant of productivity performance, rather than species richness per se. However, in reality many Mediterranean grasslands represent secondary successional communities, which are dominated by annual grasses, especially at the earliest stages if herbivory and grazing occur. Perennials may arrive at later stages and can become dominant through their more efficient use of water and nutrients. At this point their presence may mask or overwhelm subtle relationships between the diversity of more minor species and primary productivity. In fact, annuals and perennials can coexist in mosaic patches, as described in the example of Stipa-dominated grasslands (in the description of grassland communities above) where their coexistence is believed to confer stability (Clark et al. 1998). Manipulation experiments have also examined the influence of other measures of functional diversity, such as the presence of nitrogen-fixing legumes. CrossEuropean grassland experiments have shown that the presence of leguminous plants increases nitrogen accumulation, though the results were weakest for Greek and Portuguese sites compared with others (Spehn et al. 2000), possibly due to low availability of phosphorus which is necessary for adequate root nodule development. Where phosphorus and other nutrients are freely available, the presence of legumes generally increases nitrogen availability to non-legumes, although this depends on the legumes present as they differ in their efficiency of nitrogen-fixation. At some of the sites, legumes performed a key productivity role, thus supporting their identification as keystone species. In Mediterranean ecosystems, they commonly occur with grasses, with mutualistic associations developing. Together they are especially abundant in the early and middle successional stages (Blondel and Aronson 1999). Many leguminous taxa are members of the Fabaceae (pea) family, which contributes an important component to Mediterranean floral species richness. Globally, Fabaceae contains some 8,000 taxa in over 600 genera; a large proportion, though not all, are nitrogenfixing herbaceous and shrubby legumes, both annual and perennial. These are often present in early postdisturbance communities and their contribution to four regional floras of the Mediterranean has been assessed (Blondel and Aronson 1995). In the more arid regions, such as Tunisia and Israel, they accounted for 12.5 and 14.9 per cent of the regions’ Mediterranean elements respectively, compared with 9 per cent in both of the more humid regions of the Hérault (France) and Catalonia (Spain). The herbaceous perennial legumes were especially abundant.

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There is a danger in extrapolating from field experiments in one or a few types of ecosystem to other ecosystems, but the results begin to support the idea that more diverse ecosystems are more productive and more stable. However, the mechanisms by which species coexist and maintain diversity are still the subject of much debate, and despite the abundance of hypotheses there has been no clear demonstration of actual maintenance processes in species-rich ecosystems (Tilman 2000). In the context of the Mediterranean region, a possible hypothesis as to how this operates relates to the existence of spatially heterogeneous habitats in which more species can coexist because of niche differentiation, because species differ in their resource use. Habitat differences, due to steep environmental gradients in factors such as microclimate, relief, soil moisture availability, etc., mean that each species performs best in only a portion of sites. As habitat or niche heterogeneity increases further, so too does species diversity leading to greater efficiency of resource capture and use; diversity increases the likelihood that species that are better able to exploit conditions are present, and the actions of different species in different niches complement one another. From this it might follow that lower species diversity could reduce plant productivity through declining niche complementarity and less efficient use of resources. This suggests that the variety of habitat and landscape elements in the Mediterranean region promotes species richness, which in turn promotes ecosystem stability (see Zamora et al. 2007). Of course, a range of interrelated factors influences ecosystem functioning. In addition to biotic composition, i.e. vegetation communities and their associated animals, there are abiotic factors, for example climate and soil type, and a variety of disturbance regimes. Historically, Mediterranean habitats and landscapes have been long subjected to disturbances, such as fire, grazing, and a range of human activities, which should therefore be regarded as positive factors in maintaining stability. This is not to argue that human activity does not result in landscape degradation, more that over the millennia of the Holocene it has done much to promote it (Grove and Rackham 2001). Modern trends, however, may be polarizing towards more intensive activity in some places, but with abandonment of landscapes in others. Increasingly global changes, such as climate change, elevated concentrations of atmospheric carbon dioxide, nitrogen deposition, and, potentially, changing levels of ozone within both the troposphere and stratosphere are also important. The consequences of all these changes might be increasingly homogeneous landscapes, with reduced species richness and potentially, therefore, reduced ecosystem stability.

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The Role of Disturbance in Mediterranean Ecosystems Some of the highest levels of species richness anywhere in the world are found in moderately grazed shrublands and woodlands of the eastern Mediterranean. This applies at the level of alpha diversity, that is the number of species found within local assemblages or communities (Naveh and Whittaker 1979), and is very likely to be associated with repeated disturbance events such as fire, grazing, and felling, and with the incidence of drought in a semi-arid climate. This has led to recognition of perturbation-dependent ecosystems: maintenance of ecosystems requires perturbations, such as fire and grazing, to survive. Post-disturbance vegetation dynamics return communities to their pre-disturbance state, maintaining diversity (Naveh 1994, 1995). If disturbance ceases, then diversity declines with a reduction in the efficiency of ecosystem functioning. From the ideas outlined in the previous section, this might lead to a reduction in productivity as well as a reduction in the resilience of such ecosystems. The role of perturbation in ecosystem functioning is believed to operate at small scales, such as the disturbances caused by ant hills, herbivory by snails and insects, holes dug by small mammals, as well as at the larger communitylevel scale (Blondel and Aronson 1999). Thus perturbation appears to be a principal factor explaining the spatial heterogeneity of Mediterranean landscapes (Chapters 5 and 23).

Fire as a Disturbance Factor The identification of perturbation-dependent ecosystems has led to a revision in thinking about the role of fire (see also Chapter 19). Fires burning out-of-control in the Mediterranean region are regular news items during summer months and may, of course, be economically devastating and lead to loss of life. Prior to the

1980s it was estimated that as much as 10 per cent of all forests and shrublands burnt annually (di Castri 1981). These figures may now be higher as fire frequency appears to be increasing. Since the 1960s, the number of fires and the surface-area burnt in the European Mediterranean have increased exponentially (Pausas and Vallejo 1999). However, the inter-annual and inter-regional variability of fire occurrence is high. Fire is common in all the world’s mediterraneanclimate regions, though in the Mediterranean region is less likely to be natural, for example following lightning strikes. Instead, fire has been associated with people since at least the Neolithic (Naveh 1975) and possibly for as long as 300,000 years (Trabaud et al. 1993). In consequence, mediterranean-type ecosystems are fireadapted with species surviving, regenerating, and reproducing after fires (Table 7.2 and Figure 7.11). Although it is difficult to be certain as to whether the adaptations are to fire itself or to other disturbances, certain characteristics may be more pronounced as a result of the frequency or intensity of fire (Trabaud1987). Flammability is a measure of the ability of a fuel type to ignite and sustain a fire. Laboratory tests on the leaves of twenty-four dominant Mediterranean taxa have been used to categorize flammability (Table 7.3). Leaves are the most flammable part of a plant during the initiation and propagation of a fire due to their high surface-area to volume ratio and essential oil content, but the role of moisture is also important. The fresh foliage of Mediterranean shrubs generally becomes ignitable when moisture content drops below 75 per cent, and the potential for ignition is significantly related to moisture content of both live and dead fuels in grasses, shrubs, and forests. Moisture acts as a heat sink, as heat is used for the evaporation of moisture content. Moisture also dilutes volatiles and excludes oxygen from the area of combustion. The effect of moisture is determined, to an extent, by leaf structure. Least flammable taxa are those with adaptations to reduce water loss, or those with a parenchyma rich in cellular water in which more heat is required

TABLE 7.2. Fire-adapted strategies of some selected Mediterranean taxa Vegetative resprouters Obligate Quercus coccifera Pistacia lentiscus Ceratonia siliqua

Facultative Cistus salvifolius Salvia triloba Thymus capitatus Erica arborea Arbutus unedo Arbutus andrachne

Source: After Naveh (1974).

Stimulation of flowering

Stimulation of seed germination

Many geophytes—members of the following families: Iridaceae Lilliaceae Amaryllidaceae

Pinus halepensis Cistus monspeliensis Cistus albidus Rosmarinus officinalis Lavandula stoechas

Increased seed liberation and dispersal Pinus halepensis Pinus pinaster Pinus brutia

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Fig. 7.11. Basal regrowth of Arbutus unedo in the spring of 2004, following fire in the summer of 2003, Monchique, southern Portugal (photo: Harriet Allen).

to evaporate the fuel moisture. Taxa with high levels of essential oils are amongst the most volatile, including Eucalyptus camaldulensis. Eucalyptus is a widely planted, introduced genus of tree that exacerbates fire risk. Compared with foliage, branches of typical Mediterranean taxa have higher particle density, a lower surface-area to volume ratio, lower total and silica-free ash content, and a lower heat content. As a result branches are less flammable (Dimitrakopoulos 2001). The ‘literature on fire is replete with generalizations regarding fire effects and vegetation response. Like the politician who searches the scriptures for phrases to match an ideology, the assiduous shrubland ecologist can find data to support any plausible and some implausible generalizations’ (Christensen 1985: 86). This situation is not the result of poor research but is more likely due to the high variability of shrubland fire regimes. Nevertheless, in mediterranean-type ecosys-

tems, studies of the vegetation dynamics of post-fire communities show that they rapidly return to a state similar to that of the pre-fire community in terms of species composition (Trabaud 1994). There is a large abundance of herbaceous taxa in the first few years after a fire, and many of these go on to become dominant members as the community matures, although gradually the structure of the community becomes more complex, with division into numerous vegetation layers. Fire has an influence on soil processes as well as vegetation dynamics and interactions occur between the soils and vegetation. Christensen (1994) summarized some of the effects of fire on Mediterranean soils as follows. Loss of vegetation and soil surface litter increases the intensity of raindrop impact and the amount of precipitation that reaches the ground. Consequently, there may be greater soil erosion immediately after a fire. This can lead to loss of sediment-bound nutrients as

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TABLE 7.3. Flammability of selected Mediterranean plants based on laboratory tests of leaf ignition of Cretan species Species

Common name

Leaf characteristics

Less flammable Calicotome villosa Sarcopoterium spinosum

Spiny broom Thorny burnet

Juniperus oxycedrus Tamarix smyrnensis Castanea sativa Nerium oleander Platanus orientalis

Prickly juniper Tamarisk Sweet chestnut Oleander Plane

Thorns and spines which have high lignin content and low water permeability Extremely high silica-ash content

Moderately flammable Quercus coccifera Cistus salvifolius Cistus creticus Phlomis fruticosa Ceratonia siliqua Pistacia lentiscus

Kermes oak Sage-leaved Cistus Jerusalem sage Carob tree Lentisk

Hard leathery leaves with waxy or hairy epidermis which prevents water loss from evapotranspiration

Flammable Pinus brutia Pinus halepensis Quercus ilex Quercus pubescens Cupressus sempervirens Olea europea Erica arborea Arbutus unedo Pistacia terebinthus

Calabrian pine Aleppo pine Holm oak White oak Funeral cypress Olive Tree heath Strawberry tree Turpentine tree

Foliage has high surface area to volume ratio which facilitates water loss and heat absorption

Bay/laurel Eucalyptus

Rich in essential oils

Extremely flammable Laurus nobilis Eucalyptus camaldulensis

soil communities through erosion, unless they are too rapidly and intensively grazed by sheep and goats.

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Source: After Dimitrakopoulos and Papaioannou (2001).

topsoil is removed and to leaching of ions as rainfall infiltrates and percolates through the soils. Some nutrients, such as those containing nitrogen and sulphur, are lost through oxidation of organic compounds to gaseous form. Solid compounds may also vaporize, depending on the temperature of the fire, and nutrients may be lost further as ash blows away. However, not all nutrients are removed and those remaining contribute resources for the first plant colonizers. Many of these are nitrogenfixing legumes. Thus there is often a burst of flowering activity soon after fires, which aids the rapid regeneration of fire-prone vegetation communities. There may also be greater microbial activity and decomposition in the soils resulting in higher rates of mineralization and transformations such as nitrification, unless the fires have been so hot as to eliminate soil microbes. The occurrence of fire maintains a mosaic of communities, provided that the frequency is not too great. Fires do not necessarily have adverse effects on vegetation and

As characteristic Mediterranean disturbance factors, grazing and browsing are frequently considered alongside the effects of fire. The actions of sheep and goats (Figure 7.12) have traditionally been viewed negatively, being responsible for vegetation and soil degradation, even desertification (Margaris et al. 1996). However, this view should not be accepted uncritically. Like fire, grazing has been integral to Mediterranean communities since sheep and goats were domesticated some 10,000 years ago (Legge 1996; Uerpmann 1996). Moderate levels of grazing promote species diversity (Bergmeier 1998) and protect some open and semi-open communities from invasion by trees and woody species. Prior to 1948, grazing and woodcutting were widespread in the foothills of the Judaean Mountains of central Israel. This was followed by a period of effective control of wildfires and only limited grazing (Perevolotsky and Haimov 1992). By the mid-1980s, dense evergreen woodland had developed, composed mainly of dwarf shrubs (commonly Cistus species), moderate-sized shrubs (such as Pistacia lentiscus), and tall shrubs (dominated by Quercus calliprinos and Phillyrea latifolia up to 3–4 m in height). A consequence of the encroachment of woody species was a reduction of overall species diversity per unit area as less light was able to reach the ground flora, as well as increased accumulation of potential fuel for fires. Moderate grazing and fire are essential for maintaining species richness and landscape diversity, with vegetation and associated animal communities being the result of closely interwoven natural and cultural processes (Chapters 5 and 23).

Ecosystem Response to Introduced and Invasive Species Since the first trading routes were established between the Mediterranean and other parts of the world, exchanges of plants and animals have taken place, sometimes deliberately, other times accidentally (Groves and di Castri 1991). Many introduced taxa have become naturalized, or have economic value as crops or domesticated animals, or have aesthetic appeal. Some, however, have become pests. Mediterranean ecosystems, in general, appear to be prone to invasive taxa but the Mediterranean basin is less affected than other regions such as the Cape Floristic Kingdom (Cowling

Fig. 7.12. A goat browsing on Quercus coccifera, Psilorites, Crete (photo: Harriet Allen).

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1992). Nevertheless, there are invasive species that pose a variety of ecosystem problems. There are some suggestions that abandoned fields may be more prone to invasive taxa (Mooney et al. 2001) particularly by ‘weedy’ species that take advantage of the disturbed and often fertile soils. Around the Mediterranean, two successful weeds are corncockle (Agrostemma githago) and knapweed (Centaurea cyanus). They spread into the Mediterranean basin from the Near East as agriculture spread—further examples of Zohary’s law on the expansion of Irano-Turanian flora (see above)—making the most of available nutrients, lack of competition, and an opportunity to set seed before being destroyed by ploughing or harvesting. These species spread in the early Holocene, and arguably should no longer be regarded as invasive, but more recent invaders are also found in abandoned fields, such as Conyza sumatrensis and C. canadensis, both erygerons, members of the Asteraceae (daisy) family. Both originate from the New World and have become widespread throughout Europe in the last 150 years. In southern France, C. sumatrensis is established and persists on old fields, while C. canadensis is restricted to recently disturbed sites (Thébaud et al. 1996). While both have similar ‘ideal’ weed characteristics, the difference in their success as invaders appears to lie in their relative ability to compete with other species in their target communities. C. canadensis allocates more resources to the production of flowers and seeds than C. sumatrensis, but has a lower competitive ability. Hence C. canadensis is more successful invading newly disturbed sites where seedlings do not compete for scarce resources. Some taxa are able to out-compete elements of the native flora which could lead to a homogenization of vegetation communities. An example would be the hottentot fig (Carpobrotus edulis), which originates from South Africa. It is a succulent creeper that can extend over large areas of dry coastal rocks, as in the western Mediterranean, where it out-competes the scattered native species (Mabberley and Placito 1993). A potential danger of invasive species is that, if successful in out-competing native flora, there will be an overall impoverishment and homogenization of vegetation communities and ecosystems, which might reduce their productivity and stability. However, it should not be assumed that all introduced species are potentially invasive. As demonstrated by the example of Conyza sumatrensis and C. canadensis, subtle biological differences are crucial in determining the successful establishment and eventual distribution of invasive taxa, which renders generalizations difficult to make. However, the degree of invasiveness may alter with global change.

The prickly-pear cactus (Opuntia ficus-indica) is a New World introduction. In the Iberian Peninsula it is invading old fields and there is a possibility that its expansion may be favoured by the predicted further increases in atmospheric carbon dioxide (Nobel and Decortazar 1990).

The Origin of Sclerophyllous Ecosystems The origin of sclerophyllous shrublands in the Mediterranean region is a contested issue (Allen 2003), with the debate polarizing between their spread through the region as a response to Holocene climatic changes and the actions of people. Reviewing the evidence for Holocene climate changes, Roberts et al. (2001a ) note that for the circumMediterranean region, there is a complex rather than simple pattern of change, potentially due to longitudinal shifts in atmospheric circulation. In other words, climate was not uniform across the entire area throughout any particular period, and nor were climate changes synchronous (see also Chapters 3, 4, and 9). Nevertheless they recognize marked climatic differences between the two halves of the Holocene. Pollen records from a number of sites across the region indicate the presence of deciduous, often oak-dominated, taxa prior to 6,000 (uncalibrated) years BP, suggesting wetter conditions than in the second half of the Holocene: for example at sites in the Iberian Peninsula (e.g. Sierra de Cebollera (Garcia et al. 2002)), in the northern Apennines of Italy (Lowe and Watson 1993), in Sicily (Sadori and Narcisi 2001), and in the eastern Mediterranean, such as central Anatolia in Turkey (Roberts et al. 2001b). However, climatically driven interpretations of the pollen record are equivocal, due in part to the topographical diversity of pollen-bearing sites across the region and their suitability for the preservation of an unambiguous climatic signal (Chapter 9). By 6,000 (uncalibrated) years BP reconstructions of European biomes point to the existence of a more humid and cooler mediterranean-type climate, which accords with general circulation model simulations predicting lower winter temperatures at the time, based on lower early Holocene winter insolation values (Prentice et al. 1996). From mid-Holocene times forest cover generally declined and pollen from deciduous oak was replaced by that of evergreen taxa, though comparisons of circum-Mediterranean pollen records suggest that the establishment of sclerophyllous communities was not necessarily synchronous, and that

Vegetation and Ecosystem Dynamics

emplacement was transitional from both east to west and south to north (Figure 7.13 and Chapter 9). This has been interpreted as a response to increasingly arid conditions from the mid-Holocene onwards and establishment of the modern mediterranean-type climate. In contrast to the argument for the climatically driven establishment of sclerophyllous shrublands, there are many advocates for an anthropogenic origin following the spread of agriculture from the Near East about 10,000 years ago (see, for example, Quézel 1999; Tomaselli 1981b). An anthropogenic explanation is consistent with the representation of sclerophyllous shrublands as degraded ecosystems as advocated, for example, by Turrill (1929) for the Balkans. He regarded forest as the climatic climax of most of the Mediterranean region, with maquis communities originating through over-exploitation of forest with further degradation to garrigue and steppe communities (Figure 7.8). Maquis as a degradation stage in succession, either progressive or retrogressive, was an influential idea in studies of Mediterranean ecosystems (for example, Tomaselli 1981b) and evergreen shrubland communities created by such disturbances have been regarded as disclimax or paraclimax communities which could revert to the climatic climax communities if disturbances ceased. The applicability of the concept of vegetation succession to mediterranean ecosystems is now being reassessed. In part this is due to the wider recognition that single pathway, unidirectional, and progressive succession is too simplistic (see e.g. Walker and del Moral 2003). For regions with a mediterraneantype climate competition for light may not be the primary driver of supposed successional changes; moisture stress is more likely to be a limiting factor. A tree might not have a competitive advantage over lower-growing shrubs (Blumler 1993). Thus woodland cover might not be the community most suited to the varied semi-arid climates of the Mediterranean. Detailed studies of modern woodlands show their structure and composition to be heavily influenced by human activities. Closed-canopy holm oak (Quercus ilex) forests, such as those found in Spain and southern France, have been managed since at least Neolithic times, generally as coppice plots (Terradas 1999). Charcoal production began in the Iron Age and was widespread until the nineteenth century. By the twentieth century it had mostly disappeared, except in a few localities in Portugal, North Africa, and the eastern Mediterranean (Blondel and Aronson 1999). Continued, and sustainable, exploitation of holm oak forests is believed to have promoted their survival and led to homogeniza-

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tion of the woodlands at the expense of less-resistant taxa. Today, most forests bear evidence of cultivation and coppicing (Terradas 1999). Without intervention, closed-canopy holm oak forests are believed to be incapable of self-regeneration: few oak saplings grow in the understorey of established forest, despite the presence of seedlings. It seems likely that recruitment of seedlings to the sapling stage fails because they are not disturbancetolerant. Fire eliminates most seedlings and acorns, but does not kill the established trees as these are able to resprout (Retana et al. 1999). Production of acorns and their germination is then delayed until the resprouted individuals reach reproductive age. Resprouting individuals also out-compete seedlings that survive the fire, because of their already-existing root systems. A consequence of the millennial scale management of holm oak forests and occurrence of fire, is that new genotypes are not introduced into forest communities, causing gradual senescence and eventual degradation. However, in southern France there is, as yet, no sign of decline, despite continued existence since the Middle Ages.

Mediterranean Ecosystems in The Future Response to Elevated Levels of Atmospheric CO2 Contemporary climate changes are likely to have an effect on ecosystem dynamics in the Mediterranean. Increased levels of atmospheric carbon dioxide mean that a moderate increase in net primary productivity might be expected assuming plants are able to make more efficient use of available water (Mooney et al. 2001): experiments, under elevated concentrations of carbon dioxide, indicate that many species show increased efficiency of water use (Woodward et al. 1991). Taxa most likely to benefit are those which are late-growing species, shrubs, and leguminous plants that can make most use of increased availability of carbon dioxide, in other words, drought-tolerant nitrogenfixing species. Differential species response makes it difficult to assess the effects of elevated carbon dioxide. Not all plants will be able to take full advantage of increased carbon availability; early results from Californian mediterranean ecosystems tentatively suggested that in periodically stressful, low resource environments, the increase in net primary productivity might not be that great (Harley 1995). Experiments on monoliths of Mediterranean grasslands exposed to elevated carbon dioxide also

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showed only a small increase in net primary productivity, probably as a result of low nutrient availability in the soils (Navas et al. 1995). In addition, while the initial effects of increased carbon dioxide are at the leaf level, with increased photosynthesis and decreased stomatal conductance, carbon allocation to other parts of a plant is poorly understood. There are also problems with scaling results from controlled environments, such as growth chambers, to ecosystems. Thus the precise effects of increased carbon dioxide on the functioning of Mediterranean ecosystems remain largely speculative.

Ecosystem Dynamics and Land Use Changes Divergent trends in land use changes are apparent in North Africa and the Near East and in the European Mediterranean (Barbero et al. 1990). In North Africa, population pressures are leading to major declines in forest cover, predominantly associated with collection of firewood, grazing, frequent use of fire, and expansion of cropland. Between 1965 and 1976 forest and shrubland decreased in area by 3 per cent while agricultural land increased by 5 per cent (Le Houérou 1981). By con-

trast, in southern Europe since the 1950s, rural depopulation and reorganization of agriculture following the availability of subsidies from the European Union, has resulted in recovery of forest land. In western Crete, for example, aerial photographs show an increase in forest cover of 75 per cent between 1945 and 1989 (Papanastasis and Kazaklis 1998), while in Mediterranean France, forest area increased by some 22 per cent between 1965 and 1976 (Le Houérou 1981). Among the areas most affected by land use changes are the uplands, at the economic margins of cultivation. Mention has already been made of the abandonment of terraces and the spread of maquis-type vegetation. Where the maquis taxa include especially flammable species, this could increase the risk of fire. It is feared that land use changes, such as those in the northern Mediterranean, will result in the homogenization of landscapes and an associated reduction in biodiversity, both in terms of species richness and functional attributes. Scenarios modelled for 2100 across global biomes suggest that Mediterranean biomes will experience a large loss of biodiversity because of their sensitivity to a range of factors affecting global change, especially land use changes (Sala et al. 2000). As argued

Vegetation and Ecosystem Dynamics

above, this has implications for ecosystem stability. Vegetation and ecosystem dynamics, together with soils, are the determinants of landscape heterogeneity and are therefore a vital component of habitats for natural history and recreation. For their long-term existence, Mediterranean ecosystems require continued disturbance, in the form of fires, grazing, and semi-traditional agricultural practices. Conservation measures should aim to maintain these, perhaps through extension of EU compensatory payments to farmers who follow environmentally beneficial practices and through the formal recognition of biodiversity-rich landscapes, such as the dehesas and montados of the Iberian Peninsula.

Acknowledgements The author would like to thank Roland Randall (Girton College, Cambridge), William Fletcher (University of Bordeaux) and Ian Lawson (University of Leeds) for their constructive comments, Ian Agnew (Deptertment of Geography, University of Cambridge) for the artwork, and Jamie Woodward for his encouragement.

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(1995), Conservation, restoration and research priorities for Mediterranean uplands threatened by climate change, in J. M. Moreno and W. C. Oechel (eds.), Global Change and Mediterranean-Type Ecosystems. Springer, New York, 482–507. and Whittaker, R. H. (1979), Structural and floristic diversity of shrublands and woodlands in northern Israel and other Mediterranean areas, Vegetatio, 41: 171–90. Nobel, P. S. and Decortazar, V. G. (1991), Growth and predicted productivity of Opuntia ficus-indica for current and elevated carbon dioxide, Agronomy Journal 83: 224–30. Ojeda, F., Arroyo, J., and Marañon, T. (1995), Biodiversity components and conservation of Mediterranean heathlands in Southern Spain, Biological Conservation 72: 61–72. Papanastasis, V. P. and Kazaklis, A. (1998), Land-Use changes and conflicts in the Mediterranean-type ecosystems of Western Crete, in P. W. Rundel, G. Montenegro, and F. M. Jaksic (eds.), Landscape Disturbance and Biodiversity in MediterraneanType Ecosystems. Springer, Berlin, 141–54. Pausas, J. G. and Vallejo, V. R. (1999), The role of fire in European Mediterranean ecosystems, in E. Chuvieco (ed.), Remote Sensing of Large Wildfires in the European Mediterranean Basin. Springer, Berlin, 3–16. Perevolotsky, A. and Haimov, Y (1992), The effect of thinning and goat browsing of the structure and development of Mediterranean woodland in Israel, Forest Ecology and Management, 49: 61–74. Phitos, D., Strid, A., Snogerup, S., and Greuter, W. (eds.) (1995), The Red Data Book of Rare and Threatened Plants in Greece. WWF, Athens. Polunin, O. and Smythies, B. E. (1973), Flowers of South-West Europe. Oxford University Press, Oxford. Prentice, I. C., Guiot, J., Huntley, B., Jolly, D., and Cheddadi, R. (1996), Reconstructing biomes from Palaeoecological data: a general method and its application to european pollen data at 0 and 6 ka, Climate Dynamics 12: 185–94. Quézel, P. (1981a ), The study of plant groupings in the countries surrounding the Mediterranean: some methodological aspects, in F. di Castri, D. W. Goodall, and R. Specht (eds.), Ecosystems of the World, xi. Mediterranean-Type Shrublands. Elsevier, Amsterdam, 87–93. (1981b), Floristic composition and phytosociological structure of sclerophyllous matorral around the Mediterranean, in F. di Castri, D. W. Goodall, and R. Specht (eds.), Ecosystems of the World, xi. Mediterranean-Type Shrublands. Elsevier, Amsterdam, 107–21. (1985), Definition of the Mediterranean region and the origin of its flora, in C. Gómez-Campo (ed.), Plant Conservation in the Mediterranean Area. Kluwer, Dordrecht, 8–24. (1999), Les grandes structures de végétation en région méditerranéenne: facteurs déterminants dans leur mise en place post-glaciare, Geobios 32: 19–32. Rackham, O. and Moody. J. (1996), The Making of the Cretan Landscape. Manchester University Press, Manchester. Raven, P. H. (1973), The evolution of Mediterranean floras, in F. di Castri and H. A. Mooney (eds.), Mediterranean Type Ecosystems: Origin and Structure. Springer, Berlin, 213–24. Retana J., Espelta, J. M., Gracia, M., and Riba, M. (1999), Seedling recruitment, in F. Rodà, J. Retana, C. A. Gracia, and J. Bellot (eds.), Ecology of Mediterranean Evergreen Oak Forests. Springer, Berlin, 89–103. Roberts, C. N. (1998), The Holocene: An Environmental History. Blackwell, Oxford.

Vegetation and Ecosystem Dynamics Roberts, N., Meadows, M. E., and Dodson, J. R. (2001a ), The history of Mediterranean-type environments: climate, culture and landscape, The Holocene 11: 631–4. Reed, J. M., Leng, M. J., and 8 others. (2001b), The tempo of Holocene climatic change in the Eastern Mediterranean region: new high-resolution crater-lake sediment data from Central Turkey, The Holocene 11: 721–36. Ross, J. D. and Sombrero, C. (1991), Environmental control of essential oil production in Mediterranean plants, in J. B. Haborne and F. A. Tomas-Barberan (eds.), Ecological Chemistry and Biochemistry of Plant Terpinoids. Oxford University Press, Oxford, 83–94. Ruíz, M. and Ruíz, J. P. (1986), Ecological history of transhumance in Spain, Biological Conservation 37: 73–86. Sadori, L. and Narcisi, B. (2001), The postglacial record of environmental history from Lago di Pergusa, Sicily, The Holocene 11: 655–72. Sala, O. E., Chapin III, F. S., and 17 others (2000), Global biodiversity scenarios for the year 2100, Science 287: 1770–4. Schulze, E.-D. and Mooney, H. A. (1994), Ecosystem function of biodiversity: a summary, in E.-D. Schulze and H. A. Mooney (eds.), Biodiversity and Ecosystem Function. Springer, Berlin. Sfikas, G. (2002), Wildflowers of Crete. Efstathiadis, Athens. Specht, R. (1979), Ecosystems of the World, ix. Heathland and Related Shrublands. Elsevier, Amsterdam. Spehn, E. M., Scherer-Lorenzen, M., Schmid, B., Hector, A., Caldeira, M. C., Dimitrakopoulos, P. G., Finn, J. A., Jumpponen, A., O’Donovan, G., Pereira, J. S., Schulze, E.-D., Troumbis, A. Y., and Körner, C. (2000), The role of legumes as a component of biodiversity in a cross-European study of grassland biomass nitrogen, Oikos 98: 205–18. Suc, J.-P. (1984), Origin and evolution of the Mediterranean vegetation and climate in Europe, Nature 307: 429–32. Terradas, J. (1999), Holm oak and holm oak forests: an introduction, in F. Rodà, J. Retana, C. A. Gracia, and Bellot, J. (eds.), Ecology of Mediterranean Evergreen Oak Forests. Springer, Berlin, 3–14. Thébaud, C., Finzi, A., Affre, L., Debussche, M., and Escarre, J. (1996), Assessing why two introduced conyza differ in their ability to invade Mediterranean old fields, Ecology 77: 791– 804. Thompson, J. D. (1999), Population differentiation in Mediterranean plants: insights into colonization history and evolution and conservation of endemic species, Heredity 82: 229–36. Thornes, J. B. (1998), Results and prospect, in P. Mairota, J. B. Thornes, and N. Geeson (eds.), Atlas of Mediterranean Environments in Europe: The Desertification Context. John Wiley, Chichester, 162–3. Tilman, D. (2000), Overview: Causes, Consequences and Ethics of Biodiversity, Nature 405: 208–11. Tomaselli, R. (1981a ), Main physiognomic types and geographic distribution of shrub systems related to Mediterranean

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This chapter should be cited as follows Allen, H D. (2009), Vegetation and ecosystem dynamics, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 203–227.

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8

Hydrology, River Regimes, and Sediment Yield John Thornes, Francisco López-Bermúdez, and Jamie Woodward

Introduction In comparison to the rest of Europe, Africa, and Asia, most rivers arising and flowing within the Mediterranean watershed typically drain small catchments with mountainous headwaters (Figure 8.1). The hydrology of Mediterranean catchments is strongly influenced by the seasonal distribution of precipitation, catchment geology, vegetation type and extent, and the geomorphology of the slope and channel systems. It is important to appreciate, as the preceding chapters have shown, that the area draining to the Mediterranean Sea is large and enormously variable in terms of the key controls on catchment hydrology outlined above, and it is therefore not possible to define, in hydrological terms, a strict single Mediterranean river type. However, river regimes across the basin do have a marked seasonality that is largely controlled by the climate system (Chapter 3) and, in most basins, the dominant flows occur in winter—but autumn and spring runoff is also important in many areas. These patterns reflect the general water balance of the basin as a whole, but there are key geographical patterns in catchment hydrology and sediment yield and a marked contrast is evident between the more humid north and the semi-arid south and east (Struglia et al. 2004; Chapter 21). Also, because of the long history of vegetation and hillslope modification by human activity and the more recent and widespread implementation of water resource management projects, there are almost no natural river regimes in the Mediterranean region, especially in the middle and lower reaches of river catchments (Cudennec et al. 2007).

Runoff generation on hillslopes in the Mediterranean is very closely related to rainfall intensities and land surface properties as discussed in Chapter 6. While this is probably true of most catchments, runoff generation in the Mediterranean is very sensitive to vegetation cover because of the seasonal dynamics of rainfall and the role played by extreme events. The cumulative effect of these characteristics is a specific set of management problems and restoration issues and, although these are rather different in the various socio-political regimes of the region, it can be argued that they are in many ways unique to Mediterranean catchments.

The Water Balance and Mediterranean River Regimes The Annual Water Budget The broad water balance for the Mediterranean basin as a whole is summarized in Figure 8.2 which is based on the work of Jean Margat (Chapter 21) and published in Grenon and Batisse (1989). This shows mean values for annual water fluxes in the region in km3 (1 cubic kilometre = 1 billion m3 ). It is important to realize that these values will vary from year to year and they mask very significant spatial variations in catchment hydrology as we explain below. Nonetheless Figure 8.2 provides a very useful illustration of the magnitude of the annual water exchanges in the region between the atmosphere, the terrestrial catchment systems, and the Mediterranean Sea itself. Annual precipitation input to the Mediterranean ‘catchment’ is of the order of

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Fig. 8.1. An upland river catchment in the mountains of north-west Greece (photo: Jamie Woodward).

1,070 km3 each year with total evaporation losses of around 685 km3 per year. Taking into account contributions from larger river systems (such as the Nile and the Rhône that drain areas beyond the Mediterranean region), groundwater flows, and human water uses, this

budget gives an annual terrestrial water yield to the Mediterranean Sea of about 477 km3 . This component of the water balance in relation to the annual exchange of water between the Mediterranean Sea and the Atlantic Ocean is discussed in Chapter 2 and all components of

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Fig. 8.2. The water balance of the Mediterranean region showing the major fluxes (in km3 per year) between the main components of the hydrological cycle (modified from Grenon and Batisse 1989, and based on the work of Jean Margat). Values in brackets represent earlier estimates of these components.

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Fig. 8.3. Total annual runoff from river basins (shaded area) in each country bordering the Mediterranean Sea. The runoff values are given in km3 per year. The runoff volumes from river catchments in Libya, Israel, and Palestine (shown as Â) are negligible at this scale. The value given here for the Nile reflects the impact of the High Aswan Dam (modified from Grenon and Batisse 1989).

this water budget are expected to change significantly in coming decades given the predictions for climate change in the region presented in Chapter 3. An examination of annual catchment runoff to the Mediterranean Sea on a country by country basis shows large spatial variations in river discharge. Figure 8.3 shows the average annual flows (km3 ) as they were in 1985 and reported by Grenon and Batisse (1989) as part of the Blue Plan initiative. The main feature here is the marked contrast between the perennially flowing rivers on the north side of the basin and the very small contributions made by the strongly seasonal and ephemeral river catchments on the south side of the basin and in the Levant. In fact, the total contribution from river systems in Morocco, Algeria, and Tunisia (12.6 km3 per year) is almost matched by the combined runoff from just Corsica and Sardinia. Fluvial runoff to the Mediterranean Sea is dominated by river systems in Italy, Turkey, France, Greece, and the countries of the former Yugoslavia. Flows from the Nile have fallen dramatically since the construction of the High Aswan Dam and major irrigation schemes (Woodward et al. 2007).

The total annual flow of Mediterranean rivers depends on the annual water balance partition in their catchments. This has been studied and modelled by Eagleson (1978) and is depicted in Figure 8.4. This shows how annual rainfall input is divided into evapotranspiration loss from soil moisture, seepage loss to groundwater, and rainfall excess (that becomes river runoff). As mean annual precipitation increases from left to right along the horizontal axis, the rainfall excess increases, as shown in the shaded area. Thus the vertical line (1) represents the water-balance of rivers from the semi-arid and arid parts of the Mediterranean, such as those of south-east Spain, North Africa, and Israel. Vertical line (2) is the lower boundary of the sub-humid water balance regimes, found on the northern edges of the Mediterranean basin where rivers have sufficient water to flow the year round, and the soil moisture is sufficient to support a ‘good’ ground cover of vegetation. Finally, line (3) represents the ‘humid’ Mediterranean regimes of rivers with wooded mountain catchments (Chapters 3 and 7) and, of course, there are distinctive hydrological balances for special cases such as wetland systems (Chapter 9) and karst landscapes (Chapter 10).

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Fig. 8.4. Eagleson’s decomposition of the annual water balance for catchments in different climatic settings. The shaded area is the ‘rainfall excess’ or the runoff to rivers. See text for further explanation (modified from Eagleson 1978). The groundwater runoff component is especially important in many Mediterranean catchments because of the widespread occurrence and thickness of limestone rocks (Chapter 10).

Seasonal Distribution of River Flows Although all Mediterranean rivers have flows that have been greatly affected by human activity, the monthly pattern of flows or river regimes still reflects the broadly biophysical controls of climate and to a lesser extent groundwater storage. These climatic effects produce the major spatial distributions of vegetation types (Chapter 7) which combine to form key controls on land degradation as discussed in Chapter 20. The seasonal distribution of river flows is controlled mainly by the seasonal distribution of rainfall as broadly shown in Figure 8.5. This diagram also provides an approximation of the range of river regime types across the Mediterranean region. The great allogenic rivers such as the Rhône and the Nile have headwaters outside the Mediterranean winter rainfall regime and have hybrid regimes and the Rhône is considered in more detail below. Typically the Mediterranean has rains in winter and autumn (Chapter 3) with the middle and lower reaches of many rivers being dry in the summer (Figure 8.6). This is characteristic of much of the Maghreb, the Levant, southern Greece, southern Italy, the very south-east and south of the Iberian Peninsula and most of Portugal. Most of the Iberian Peninsula, Italy, and the western parts of the Balkan Peninsula also receive significant autumn rains. Only the continental parts of the Balkan states and Macedonia have rainfall

all year round. Monthly flow series for eight river gauging stations across the Mediterranean region (and one in the headwaters of the Rhône) are given in Figure 8.7. The histograms show the range of river regimes evident in the Mediterranean moving clockwise (a) to (i) around the basin (Figure 8.5) and include river systems in France, Italy, Albania, Turkey, Cyprus, Israel, Algeria, and Morocco. All the stations (b) to (i) show a marked summer decrease in river flows with the channel bed of the Vasilikos River (151 km2 ) in Cyprus drying out completely for five months from July to November. The flow data for the rivers in the east and south of the Mediterranean (e) to (i) highlight the importance of winter and spring precipitation in these catchments. The Rhône at Chancy, on the French–Swiss border, has a mean annual discharge of 342 m3 s−1 . The minimum and peak recorded flows at this station in the period January 1965 to December 1982 were 134 and 740 m3 s−1 respectively. This regime reflects the essentially Alpine character of the headwaters, with a strong spring and summer snowmelt. By the time the river reaches Beaucaire in the Garrigue, the Mediterranean regime is already evident. The catchment area here is almost 100,000 km2 , about nine times that at Chancy. The mean discharge is >1,600 m3 s−1 with minimum and maximum flows of 420 and 5,077 m3 s−1 respectively. The strong dip

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Fig. 8.5. Rainfall regimes in the Mediterranean region (modified from Huttary 1955). The solid line is the watershed of rivers that drain to the Mediterranean Sea. The letters (a to i) mark the locations of the gauging stations in Figure 8.7.

in the summer and early autumn months (July to October) is quite characteristic of the Mediterranean regime. Snowmelt makes an important contribution to many headwater river systems in the Mediterranean region—especially where the mountains rise above 2,000 m (Chapter 12). The regime for Wadi Moulouya in Morocco—at the Dar el Caid station—shows a more pronounced Mediterranean regime. Here the annual mean discharge is 20 m3 s−1 with minimum and maximum peak flows of 0 and 327 m3 s−1 respectively. August has the lowest average flow (3.6 m3 s−1 ) and April, the maximum, showing the spring rainfall effect (Figure 8.7). The catchment of the Wadi Moulouya at Dar el Caid (24,422 km2 ) is similar in size to the Menderes River at Soke (23,889 km2 ) in Turkey yet, due to the greater aridity of the local climate, its mean annual discharge is only about one fifth of the Menderes (Figure 8.7). For large parts of the semi-arid Mediterranean, channels may be without flow for most of the summer but commonly restarting with the onset of autumn storms. Surface channel flows in these rivers normally cease in early May, though flow will often continue beneath the gravel beds and may be used for irrigation throughout the summer months. Many channels are truly

ephemeral and convey flows only during major storm events (Cudennec et al. 2007)). In flash floods, water drains into the channel beds so the flows may only be evident in upland bedrock reaches and then disappear in the middle and downstream reaches where the channels are formed in thick coarse-grained gravely sediments (Thornes 1977; Butcher and Thornes 1978; Shannon et al. 2002). A key problem in many parts of the semiarid Mediterranean is the absence of reliable, long-term records of river flows.

Human Modification of Mediterranean Catchment Hydrology The human history of the Mediterranean lands has been strongly moulded by its rivers (Lewin et al. 1995; Grove and Rackham 2003; Chapter 11) and among all the natural ecosystems, rivers and floodplains are the most intensively used by Mediterranean people. The floodplains have formed important corridors for the migration of plants, animals, and people and have also provided some of the most fertile alluvial soils in the

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Fig. 8.6. The seasonally dry gravel-bed channel of the Voidomatis River upstream of the Vikos Gorge in north-west Greece (photo: Jamie Woodward). Note the high flow stage marks on the bedrock wall.

world, on which intensive irrigated agriculture has been established. Around them, complex legal systems for their use have developed. As in other parts of the world, the human modification of river flows and channels has taken place both directly and indirectly. Indirect changes are those which change runoff generation and sediment supply at the basin scale—mainly through changes in land use. Direct changes are those which alter the flow regimes, sediment stores, and planform of the channel, usually by engineering works related to water resource development and flood protection or alluvial gravel extraction (e.g. Christopoulos 1998; Nicholas et al. 1999; Kapsimalis et al. 2005) (Chapter 11). What sets the Mediterranean region apart from other areas of the world, however, is the long history and widespread occurrence of both direct and indirect modifications to catchment hydrology and river channel systems (VitaFinzi 1978; Macklin et al. 1995).

Indirect Changes to Catchment Systems Vegetation Cover and Runoff Dynamics Elwell and Stocking (1976) showed that runoff from experimental plots was reduced as the vegetation cover was increased and the curve is exponential (Figure 8.8). On the horizontal axis is the percentage vegetation cover in plan view. On the vertical axis is the amount of runoff as a percentage of bare soil. This parameter is the result of many complex factors, as discussed later in this chapter, most prominently rainfall intensity, basin morphology, and soil characteristics, such as infiltration capacity and crusting. These experiments explain the earlier discoveries of American workers and the speculations of the ancient Greeks (Grove and Rackham 2003) and have been reconfirmed in many studies since, in

(a) Rhône at Chancy, Switzerland Catchment = 10,299 km2 Mean discharge = 341.61 m3 s-1

(b) Rhône at Beaucaire, France Catchment = 95,590 km2 Mean discharge = 1,694.95 m3 s-1

(c) Tiber at Rome, Italy Catchment = 16,545 km2 Mean discharge = 231.13 m3 s-1

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Fig. 8.7. Average monthly flows for rivers around the Mediterranean basin based on data from various sources cited in the text. Catchment sizes vary greatly, but note that the water fluxes for rivers draining the northern part of the basin, (a) to (e), are at least an order of magnitude higher that those draining the east and south of the basin, (f) to (i).

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theoretical (Carson and Kirkby 1972), experimental (Francis and Thornes 1990), and field investigations (Bochet et al. 2006). In the last paper, the authors found that, for the three species of Mediterranean shrub considered, whereas total vegetation canopy and litter cover reduced soil loss and runoff rates, soil loss and runoff increased with plant volume. A very careful and detailed study of the effect of canopy structure on rainfall intensity for three common shrub species of the western Mediterranean (Anthyllis cytisoides, Stipa tenacissima, and Retama sphaerocarpa) was undertaken in laboratory simulations by GarciaOrtiz (2006). The species are shown in Figure 8.9. She found significant differences between species in canopy drip and through-fall volumes (as a percentage of the precipitation) and a correspondence between the main partitioning pathways and the root system type of the species was identified. Species with a deep root system were more efficient at redistributing rain by stem flow whereas species with a shallow root system redistribute rain mainly as canopy drip (Stipa tenacissima). Retama sphaerocarpa is the species with the least canopy storage, while Anthyllis cytisoides has the highest. The effects of rainfall intensity could not be fully explored and there is a need here for further research. It appears that Darwinian selection has led to plant morphologies that efficiently use the available rainfall to ensure their survival and may explain why these species are so abundant.

It follows that changes in the vegetation cover will transform the runoff regimes of Mediterranean catchments. This proposition is theoretically and empirically sound, but the results of studies of deforestation have not always been so clear. This is partly because of the wide variations in what is regarded as ‘forest’, because deforestation covers a wide variety of patterns of cutting within a catchment and because post-deforestation recovery depends on the sequence of weather in the years following cutting. These factors have been studied in detail by Obando (2002) in a semi-arid Mediterranean environment and the results suggest caution is needed in overenthusiastic endorsement of reafforestation as a strategy for reducing runoff. Chirino et al. (2006) concluded from their study of the effects on runoff of a thirty-year-old Aleppo pine plantation in south-east Spain that the conservation of natural communities or the restoration of semi-arid ecosystems with native shrub species can be as effective as reafforestation in managing runoff and erosion as suggested by Francis and Thornes (1994). The same care needs to be exercised when claiming major impacts on catchment runoff resulting from fire and grazing. A large number of studies have been carried out on the effects of fire on water and sediment yield in the Mediterranean (Wainwright and Thornes 2003). Most of these results demonstrate an increase of at least one order of magnitude in both runoff and sediment yield. Variations in the intensity of the burn yield different results, as shown by Inbar (1992). The impacts of fire are largely in changing the infiltration properties of soils by changing their hydrophobic properties. These changes have to be balanced against the positive benefits of both reafforestation and grazing. As with soil erosion, the issue is simply not clear cut, nor are the conclusions uniform (Chapter 20). Chapter 19 deals with the wildfire hazard more generally in the Mediterranean.

Cultivation and Catchment Hydrology There is less doubt about the impacts of cultivation on catchment hydrology, though still some scope for disagreement. Rapid runoff from hillslope and channel systems is a major catchment management issue across the Mediterranean region (Figure 8.10a). Wainwright (1996) examined the controls on runoff in the September 1992 flood in southern France in the basin of the River Ouvèze. Although the flooding was principally a result of the high rainfall intensities that occurred

Fig. 8.9. The species used by Garcia-Ortiz in her study of rainfall partitioning by different Mediterranean plants. (a) Anthyllis Cytisoides (b) Retama retamo (c) Stipa tenecissima. The Retama is about 2 m high (photos: John Thornes).

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that it is unimportant. It is that the activity is so diverse and its intensity so variable through time that it is difficult to predict the outcomes of these human modifications on catchment runoff and river regimes in Mediterranean environments. Nevertheless, because the effects of vegetation on runoff are reasonably well understood, they provide us with a tool for managing grazing in a way that optimizes runoff and sediment yield by assigning areas for grazing use (Thornes 2007).

Direct Changes to Catchment Systems Terraces and Water Harvesting

Fig. 8.10. Flood flows and erosion in Mediterranean catchments. (a) Runoff during a flash flood in Tunisia (photo: Ian Foster). (b) Deep rill erosion on sloping agricultural land in the western Peloponnese, Greece (photo: Jamie Woodward).

in the catchment, experimental and modelling evidence suggest that the severe flooding could be explained by runoff production in the agricultural areas on the lower slopes and terraces immediately upstream of the village of Vaison, as well as from the badland areas which are concentrated more in the upland areas of the catchment. He indicated that cultivation practices encouraged flow concentrations, accentuating the links between slope and channel. An experimental study, using digital simulation, by Kirkby et al. (2002) indicated how steepness of the terrain and the direction of ploughing, relative to the main slope direction, influenced the volume of runoff. Figure 8.10b shows how heavy rainfall and overland flow on bare and steeply sloping agricultural land in the Alfios River basin of the western Peloponnese can lead to the formation of deep (>60 cm) rills and the loss of valuable topsoil. As with deforestation and cultivation, the argument about the impacts of grazing on catchment runoff is not

One of the most ancient and widely practised direct water control measures in the Mediterranean has been the construction of agricultural terraces (Figure 8.11) (Grove and Rackham 2003; Chapters 6 and 7). This is a skilful procedure. The terraces have to be easy to cultivate with draft animals (oxen or mules), accessible, and fairly flat, and yet designed to ensure that the water is conducted away safely without undue erosion. Water infiltrates into the terraces and, if hydraulic pressure builds up inside the terrace, it can lead to piping which in turn can produce deep shafts big enough to swallow animals and machines. In North Africa and the Levant, the management of catchments has become a fine art, in order to conserve water for agriculture (Evenari et al. 1982). This procedure is called water harvesting. In carefully selected areas, vegetation is removed and the soil ‘crusted’ by trampling or hardening. Then the runoff from these crusted surfaces is conveyed by lined ditches into tanks, cisterns (Figure 8.12), or aljibe (tanks with rounded roofs that collect evaporated water) to be used for crop irrigation. This practice was well established in Roman times and used by Arabic farmers in the Maghrebian states and Iberia (Gómez-Espin 2004). The modern equivalent of the ancient stone-fronted terraces is extensive contour benching over large areas, usually by bulldozers. These are direct conservation measures for runoff on the hillslopes and may be used for flood protection.

Check Dams and Reservoir Development Direct measures in the channels range from small check dams, a few metres high, to major reservoirs (Figure 8.13a, b). Check dams serve multiple purposes. In addition to reducing the amount and velocity of runoff from headwater channels, they also store transported

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Fig. 8.11. Stone-walled terraces on hillslopes near Campanet in central Majorca (photo: Jamie Woodward).

sediment and provide flat areas for cultivation, into which water quickly and easily infiltrates. Another small, direct control on runoff in channels that dates back to the ancient water-control era of the western Mediterranean is the boquera system. Walls are constructed in the channel floor to lead flood water into fields, or boqueras, on either side of the channel where water is stored, as soil moisture, to provide for cereal and tree crop cultivation. Mill sites also abound in channel systems and the power provided can then lift irrigation water into canal systems that may redistribute it across hundreds of hectares. Other minor direct features include tunnels constructed beneath river beds—the water infiltrates into the gravels and then into permeable brick-lined tunnels that may convey it as much as 30 km to provide town water supplies. Many inter-basin water transfers were introduced in the Mediterranean region during the Roman Period as the expansion of urban settlements increased the demand for water supplies. In southern Epirus (north-west Greece), for example, the city of Nikopolis, founded by Octavius Augustus in 30 BC, required the construction of a major aqueduct (50 km in length) that brought large volumes of water from the Louros Valley to the north to the 300,000

residents of the city (Mavromati and Chryssaidis 2007; Vita-Finzi 1978). Above all, the runoff of rivers has been modified by the great reservoirs the construction of which commenced at the end of the nineteenth century in the period of La Grande Hydraulique. Some of these have already filled with sediments (Figure 8.13b). In Murcia, Spain, the wall of the Valdeinfierno reservoir was extended vertically, while in the Puentes reservoir immediately downstream (Figure 8.14a) a new wall has been constructed in front of the original. This construction of reservoirs has had a marked impact on the flow regime of the Segura River (Figure 8.14b), completely changing it from one dominated by peak flow in winter months, to one of peak (but lower) flows in summer months. These changes reflect not only the winter rains, but also the summer demand for irrigation. The summer discharges are now controlled ‘ecological’ flows required to maintain the ecological health of the river as it passes through the city of Murcia and beyond. As a consequence of these and other measures, the quality of the Segura has notably improved since 1972. Because of the seasonal characteristics of both the industry (fruit processing) and the flows, the smaller

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Fig. 8.12. A water cistern directly under the former Greek agora in Ptolomais, Libya. It receives water via an aqueduct from mountain springs 25 km to the east (photo: courtesy of Paul Fordham).

tributaries, such as the Mula, were heavily contaminated by stagnant pools in the summer months (Victoria Jumilla and Vicente Lopez 1986). These conditions impose special demands for integrated river management under the European Water Framework Directive. Establishing an appropriate balance between water storage, water abstraction, and the ecological needs

of channel systems is a major water resource management issue across the Mediterranean (Thornes and Rowntree 2006). The loss of water storage capacity to sedimentation is a particularly important water resource management issue in Mediterranean northwest Africa. In Morocco, the average annual water storage loss to sedimentation in the thirty-four largest

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Fig. 8.13. (a) A newly built check dam in the Segura River basin, south-east Spain. The water and sediment will fill the area to the left of the main wall. The apron and raised hump to the right are to prevent a scour pit developing during flood events and undermining of the main wall. (b) A sediment-filled reservoir: Valdeinfierno Reservoir, Murcia, Spain. The water storage capacity is now less than 5 per cent of the original. The foreground is wave-rippled sediment and not water. The location of this reservoir is shown at Figure 8.14a (photos: John Thornes). (c) Oblique air photograph of the town of Puerto Lumbreras and the Rambla Nogalte (see Figure 8.14a) in the aftermath of the large flood in October 1973 (http://www.puerto-lumbreras.com/galeria.asp?galeria=3).

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reservoirs amounts to about 0.5 per cent (Lahlou 1988). Reservoirs in Algeria are estimated to be losing 2–3 per cent of their storage capacity each year and this equates to about 90 million m3 of water (Grenon and Batisse 1989).

Runoff Generation at the Local Scale In the previous section we looked at how the catchment runoff of Mediterranean rivers has been both indirectly and directly modified over the centuries by human activities. In this section the main natural controls of runoff generation at the local scale are considered. With the coming of remote sensing, hydrological studies of Mediterranean catchments included modelling runoff on the basis of the distribution of rainfall over the catchment and its distribution of intensity through time. It became recognized that soil conditions controlled infiltration rates and hence runoff, especially where vegetation was very scarce. A growing body of research indicated that surface crusting was also important. This crusting might be compacted soils a few millimetres thick formed under heavy rain, or chemical precipitation such as Mediterranean calcretes (Chapter 6). The crust can also be formed by lichens or other micro-organisms. These various crusts seriously impede infiltration rates and produce more excess runoff. In large catchments, these variations in surface texture and therefore runoff can be enough to control the volume and timing of runoff and therefore of stream hydrographs. Even the simple discrimination between bedrock runoff and alluvial surface runoff can be enough to improve runoff predictions from dryland catchments greatly. The term ‘hydrologically similar surfaces’ (HYSS) is used to describe surface properties that affect runoff generation. A study of the 1,400 km2 Nahael Zin catchment in southern Israel by Lange et al. (1999), using remote sensing to identify HYSS and GIS techniques to map and model runoff, removed much of the uncertainty in runoff prediction. Field data were used to specify the characteristics of the HYSS. By way of clarification of this emerging concept, Bull et al. (2000) used an improved infiltration model coupled with digital runoff simulation to examine the major controls on runoff generation at different scales. The study was based on observations of the Nogalte and Torreavilla catchments in south-east Spain. At the hillslope scales, runoff intensity was recognized by classifying the morphological evidence that ranges

from ‘no evidence of flow’ to ‘streams and gullies >1 m deep’. The different morphological runoff types are arranged in different zones on different lithologies according to the length-slope product. For the simulated storms, the storm structure (cell size and intensity) was not important in producing at-a-point runoff even over a simulated catchment of 1,000 km2 . However, at these larger scales, the basin morphology has an important control on runoff generation, because it determines water storage on the hillslope and transmission losses into the channel bed. Also at this scale, the connectivity of the channel network becomes very important. In summary, the most important control on runoff for these simulated conditions is variations in intensity through a storm. For observed data on catchment characteristics (e.g. shape and connectivity) and storm cell size, the effect of storm size did not lead to large errors, if estimates of catchment runoff were made from a single central rain gauge. Large divergence only appeared for areas of 1,000 km2 or more. The final factor of great importance is the areal distribution of crusting properties as indicated above (Chirino et al. 2006). Runoff generation in Mediterranean catchment systems is rather different from that of weakly-seasonal temperate environments, as discussed by Beven (2002). In this important contribution, Beven described our perceptions of what constitutes the behaviour of ephemeral channels: the complexity of the controlling processes, the strong non-linearity of that behaviour, the extreme heterogeneity of the runoff surfaces, and the limited storm extents in space and time (Figure 8.15 Chapter 18). He concluded (p. 62) that ‘We still have much to learn about runoff generation in semi-arid catchments.’ In particular he questioned the assumption that flow in these catchments was dominantly Hortonian overland flow (when rainfall intensities are significantly higher than infiltration capacities). Kirkby (1969) showed that in temperate environments this was rarely the case and supported Dunne and Black’s thesis that hydrological modelling must accommodate sub-surface flow, especially where flow is convergent (Dunne and Black 1970). Beven (2002) went on to review evidence for subsurface flow even in dryland systems, reaching the conclusion that, in Mediterranean catchments, this can be readily accepted as occurring under winter conditions. The main problems are choosing appropriate infiltration models and capturing the huge heterogeneity at different scales, even down to the level of the individual plant (Figure 8.9). This seems to require an alternative approach to fully distributed hydrological models or

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Fig. 8.15. The Río Aguas at Urra in the Sorbas basin, Almeria, south-east Spain. (a) The river in flood with turbulent flows and high suspended sediment concentrations (photo: courtesy of Joe Walsh) and (b) when the channel is dry in April 2006 (photo: Jamie Woodward).

a better management of the parameterization problem. An example of the former is the stochastic approach to modelling, as in the work in the Rambla de Algeciras (Murcia) by Conesa-Garcia and Alonso-Sarria (1997) or by Shannon et al. (2002) which proposes a Markov routing approach to ephemeral channels. Beven’s chapter concluded with the second approach—improving the parameterization. This may seem a far cry from Mediterranean rivers, but the continued demand for water and buildings in the coastal zone where the rivers and ramblas (ephemeral gravel-floored channels) meet the sea means that flash floods there can have disastrous effects on property, so that effective identification of flood-prone areas is a very high priority (Chapter 18). Although this can in part be achieved by historical evidence and experimental hardware models (e.g. for groundwater recharge), mathematical hydrological modelling also offers a route that can be pursued.

Catchment Sediment Yields The steep relief of Mediterranean river basins, the widespread occurrence of erodible rock types, and the heavy storm events that can generate rapid surface runoff combine to produce high suspended and bed sediment loads across the region. The erosion processes responsible have been reviewed in detail in Chapter 6. While there is little doubt that human activity is also important—and bare fields with limited soil conservation measures can lose large volumes of top soil during storm events (Figure 8.10b)—the underlying topographical, lithological, and climatic controls combine to produce an environment with much higher sediment fluxes per unit area than more temperate areas in central

and northern Europe for example (Woodward 1995; Milliman and Syvitski 1992).

Suspended Sediment Yield Dedkov and Moszherin (1992) have attempted to distinguish between the anthropogenic and natural or background sediment yield of river basins in mountain environments across the world. This is a very difficult task and the database for mountain river basins is very patchy indeed. However, they used land use type and extent to classify catchments into natural or modified states. Many would dispute this finding, but their analysis suggests that the sediment loads of mountain river systems in the Mediterranean are influenced by human action to a greater extent than those in any other climatic zone (Figure 8.16). It is, however, difficult to generalize about such a diverse region when there are so many variables to consider. Nonetheless, what does seem clear is that within the Mediterranean region, clear contrasts are evident in the magnitude of suspended sediment yields from catchments in the north and south with the latter typically an order of magnitude higher that the former (Milliman and Syvitski 1992; Woodward 1995). Table 8.1 contains data for two large river catchments in the northern Mediterranean that drain into the Aegean Sea to the west of Thessaloniki in north-east Greece. The Axios (23,747 km2 ) and Aliakmon (9,250 km2 ) rivers drain humid headwater catchments with elevations well over 2,000 m and have mean annual water fluxes of 5,030 and 2,292 million m3 per year respectively (Kapsimalis et al. 2005). These catchments—which include a range of land use types— have mean annual suspended sediment yields of 555 and 670 t km2 yr−1 , respectively, which are much lower than the typical yields for catchments in the Maghreb

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(Lahlou 1988; Woodward 1995; Woodward and Foster 1997). In a review of the sediment yield of African rivers, Walling (1984) identified the narrow strip of upland catchments across Mediterranean north-west Africa as the highest yielding river basins on the continent. Suspended sediment yields from many of these catchments range from 1,000 to 5,000 t km−2 yr−1 . In an average year, the total annual suspended sediment load from river systems in the Maghreb that drain to the Mediterranean Sea is about 100 million tonnes (Probst and Amiotte Suchet 1992), and the flood events responsible are commonly characterized by very high suspended sediment concentrations. Network expansion by gully development in soft sediments can provide an important source of suspended sediment in Mediterranean catchments (Figure 8.17a). Thick, single-event accumulations of fine-grained sediments in channel and floodplain zones are indicative of high suspended sediment concentrations and loads (Figure 8.17b). High suspended sediment loads are a feature of many Mediterranean river basins but this is often strongly controlled by catchment lithology. Figure 8.18 shows how sediment yields in Moroccan catchments are markedly reduced where resistant rocks such as quartzites and granites are important. In contrast, where

Fig. 8.16. Suspended sediment yield from river basins in mountain environments in different climate and vegetation zones based on the analysis of Dedkov and Moszherin (1992).

TABLE 8.1. Water and sediment fluxes from the Axios and Aliakmon rivers that drain into the north-west Aegean Sea Axios Catchment area (km2 ) Maximum elevation (m) Mean annual rainfall (mm) Land use (%) Forest cover Uncultivated area Arable land Urban area Wetland Water fluxes (106 m3 ) Mean annual discharge Monthly maximum discharge Monthly minimum discharge Material fluxes (106 t yr−1 ) Suspended sediment load Dissolved load Bed load

23,747 2,800 650

Total load

16.9 5.9 75.4 1.7 0.1 5,030 8,800 1,545

Aliakmon 9,250 2,200 750 19.0 40.0 39.3 1.3 0.4 2,292 4,320 662

13.2 1.7 2.6

6.2 1.2 1.3

17.5

8.7

Source: Modified from Kapsimalis et al. 2005.

shales, marls, flysch, and other erodible lithologies are extensive, sediment yields can exceed 6000 t km2 yr−1 .

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Syvitski et al. (2005) estimated that human activity (mainly through land use change in river catchments leading to soil erosion) has increased the global sediment load of river systems by 2.3 ± 0.6 billion metric tons per year whilst simultaneously reducing the flux of sediment reaching the marine environment by 1.4 ± 0.3 billion metric tons per year due to sediment retention in reservoirs. Dam construction for water storage has been a favoured response to the pronounced seasonality in Mediterranean river flows since Roman times and many of today’s catchments contain impoundments at a range of scales. In the case of river systems draining to the Mediterranean and Black Seas, Syvitski et al. (2005) have estimated that 30 per cent of the fluvial suspended sediment flux is retained behind reservoirs and this compares to a global average of 20 per cent. Interestingly, their analysis also shows that Mediterranean river systems have the lowest proportion of summer (JJA) and autumn (SON) sediment fluxes of any landmass or ocean basin at 9 and 7 per cent respectively of the annual total.

Bed Load Fluxes

Fig. 8.17. (a) Gully erosion in soft sediments in central Israel (photo: Ian Foster) (b) A veneer of fresh suspended sediments deposited within the channel zone of the Torcicoda River in central Sicily (photo: Jamie Woodward).

River monitoring studies and reservoir sediment surveys have shown that river catchments in the Maghreb region of North Africa have some of the highest suspended sediment yields in the circum-Mediterranean zone (Woodward 1995; Woodward and Foster 1997; McNeill 1992). Water storage losses to sedimentation are a major concern for water resource managers in Morocco, Algeria, and Tunisia (Chapter 21). In a recent global-scale analysis of human impacts on the flux of sediment from the land surface to the oceans,

Information on bed load transport in Mediterranean river systems is generally much patchier than the database on suspended sediment transport. However, valuable datasets are available from Israel in particular where pioneering work on the ephemeral rivers of the Negev has provided important insights into the dynamics of the region’s dryland rivers (e.g. Laronne and Reid 1993; Reid and Laronne 1995; Reid et al. 1998). Bed load transport is difficult to measure and it poses particular logistical problems in semi-arid and arid river catchments because floods are so infrequent and unpredictable. In this context, it is useful to consider some of the hydrological characteristics of dryland catchments in the Mediterranean. A well-researched example is the Nahal Eshtemoa catchment (119 km2 ) that drains part of the southern Hebron Hills and is typical of the river systems of the northern Negev Desert. This catchment receives an annual rainfall input of 220 to 350 mm within a region where the potential evaporation is commonly >2,000 mm per year. This produces large soil moisture deficits which, in combination with sparse vegetation cover and low infiltration capacities, can produce overland flow within minutes of the onset of rainfall (Reid et al. 1998). Flow duration analysis of the Nahal Eshtemoa catchment for the period 1991–5 revealed that the channel conveyed flows on only seven days per year (2% of the time) and that overbank flooding took place for only 0.03 per cent of the time (Reid et al. 1998). The number of flash floods over this period

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Fig. 8.18. Suspended sediment yield from Moroccan river catchments formed in different rock types (modified from Lahlou 1988).

ranged from two to seven per year and most were produced by convective storms. The trunk stream hydrographs for these events show very spiky behaviour with most of the floods showing very steep rising limbs and multi-peaked hydrographs. Due to the roughness of the coarse-grained channel bed, flash-flood bore velocities are typically quite slow falling within the range 1–1.5 m per second (ibid.). Such infrequent and unpredictable flow regimes are typical of the eastern and southern Mediterranean and contrast markedly with the more humid regions in the north (Figure 8.7). The database on bed load flux in the ephemeral river systems of the Negev shows very high rates of sediment flux in comparison to perennial rivers of more humid environments. Laronne and Reid (1993) have argued that bed load data for the Nahal Yatir catchment show it to be 400 times more efficient at transporting coarse sediment than some of its perennial counterparts in more humid zones. Annual bed load sediment flux in the neighbouring Nahal Eshtemoa system was calculated as 4,641 tonnes or 39 tonnes km−2 while suspended sediment yield over the same period was 433 tonnes km−2 yr−1 giving a bed load to suspended load ratio of 1:11 (Reid et al. 1998). Data on bed load transport are now also available for perennial river systems in the mountains of the Mediterranean basin. For example, Garcia et al. (2000) employed an automated bed load monitoring system in a small catchment (35 km2 ) in the upper reaches of the Todera River in the coastal mountains of Catalonia in north-east Spain and monitored five flood events between 1995 and 1996. A key observation was the marked variability in bed load flux during the course of each flood event. Distinctive pulses of bed load transport were associated with sudden increases in flow depth and bed load response to hydraulic variables such as critical shear stress varied markedly between flood events. Much further downstream in the

lower, ephemeral, part of the same river basin, Rovira et al. (2005) assembled a fluvial sediment budget for an 11-km reach that lies just upstream of the basin outlet (894 km2 ) to the Mediterranean Sea. For the period from January 1997 to June 1999 they showed that the bulk of the total sediment load was transported as coarse bed load sediment. Around 156,000 tonnes of sediment (80% as bed load and 20% as suspended load) entered the study reach and approximately 107,000 tonnes (83% as bed load and 17% as suspended load) of sediment exited the reach and were discharged to the coastal zone and the Mediterranean Sea. Using repeat field surveys of the channel and floodplain zone in association with monitored sediment flux data, Rovira et al. (2005) estimated that 49,600 tonnes of sediment were deposited within their study reach over this period giving a mean rate of sediment accumulation on the channel bed of 6.8 mm per year. The extraction of bed load materials for aggregates and other commercial uses is widespread in the Mediterranean and this activity can result in dramatic modifications to coarse sediment budgets in river channel systems (Nicholas et al. 1999). Information on bed load transport should form part of the design of catchment management plans to ensure that such resource exploitation is sustainable in the long term. This aspect of direct river channel modification is discussed in more detail in Chapter 11.

River Catchment Management and Hazard Mitigation Floods and Flood Protection Damage and loss of life from floods is commonplace in Mediterranean river catchments (Chapters 3 and 18) and appears to have been so throughout history. For south-east Spain, Gil Olcina (1971) records the floods of

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Lorca (Murcia) in the seventeenth century and Pocklington (1986) describes flooding of the huerta (cultivated area) of Murcia in Arabic times. Chapter 11 documents flood histories for the Little Ice Age in various parts of the Mediterranean including southern France, Corisca, Crete, and north-west Greece. An example from recent times was the flash flood of Puerto Lumbreras (Figure 8.13c). On 18/19 October 1973 an intense storm crossed southern Spain, from Torremolinos in the west to Benidorm in the east. Over 250 mm of rain fell on the mountains of the Sierra Nevada in the headwaters of rivers that drain to the Mediterranean coast, the Guadalhorce, Guadalfeo, Andarax, and Almanzora. The storm created immense damage to roads, bridges, irrigation installations, and agricultural land. It caused severe loss of life in the village of La Rabita, where homes were destroyed by a mudflow emerging from a small canyon. In the town of Puerto Lumbreras (Figure 8.13c), where the weekly market was being set up in the dry river-bed, the flood wave washed away the entire market. Houses and cave dwellings and Moorish irrigation canals were destroyed by the on-rush of water. Today the scene is more tranquil. At Puerto Lumbreras, huge walls have been constructed on either side of the channel to protect the remaining town. The catchment of the Río Nogalte (Figure 8.14a) has been contour ploughed and trees have been planted. Over a hundred check dams have been installed in the steep-sided tributaries (barrancos) by the government agency responsible for natural resources. The Ministry of Public Works has reconstructed roads and bridges. Houses have been repaired, sub-alluvial galleries reinstated and the market continues to operate in the river-bed on a regular basis. The more protection that is provided, the more risk the population appears to be prepared to take. Although statistics suggest the event has a probability of occurring about 1 in 500 years, an equally disastrous storm occurred in 1982 in the Valencia area, immediately to the north. This story is repeated frequently across the Mediterranean and reflects a combination of intense storms, poor vegetation cover, and a long period of land management that encourages high and rapid runoff from mountainous terrains and impervious soils. The storm and flood hazards in the Mediterranean region are discussed in greater detail in Chapter 18.

River Catchment Management In the EU MEDALUS (Mediterranean Desertification and Land Use) Project, the role of flooding in the Guadalentin River basin (Figure 8.14a) was recognized

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as a component of the problem of desertification and a plan was devised to manage the problem through land use control in the catchment of one of its tributaries— the Rambla Nogalte (Rojo Serrano et al. 2002). This was already partly underway in the remedial work following the 1973 flood. The rainfall of the basin ranges from 200 mm per year in the lowland areas, to 800 mm in the mountain headwaters and the rainfall is characterized by high intensities and short duration. Half of the basin has slopes greater than 12 per cent (at which erosion starts) and about one-third has slopes greater than 20 per cent (the overland flow-rilling threshold). A very small area of the basin is covered with sclerophyllous forest in the mountains. At lower altitudes and over most of the basin, the conditions are too dry for oak woodland (Quercus rotundifolia), but pines (Pinus pineaster, nigra, and halepensis) can survive. At even lower altitudes, bushes of the matorral prevail. Overall, 45 per cent of the basin is covered with ‘wildland’ and only 25 per cent of the basin is covered by forest, with another 20 per cent by shrub and bush communities. Water scarcity, groundwater exploitation, and flood protection are the most important environmental issues, followed by soil erosion. In the past, mechanized afforestation techniques in this catchment have been more effective than manual ones in decreasing hillslope runoff, retaining and storing as much water and moisture as possible (Rojo Serrano et al. 2002: 310). Pinus halepensis is the fundamental afforestation species. It is a Mediterranean species and its resistance to excessive sunshine means it has no need for shade when first planted. However, in the worst climatic and edaphic situations, this species is unable to reach forest status. In this study, specific management techniques were recommended for the Upper Guadalentin, ranging from better management of the existing forest and grazing areas, to restoration of saline soils in the lowland areas. This example provides a model approach that could be applied in many flood-prone areas of the Mediterranean and beyond. In other management schemes, unexpected effects may lead to unexpected and undesirable results. Borel (1994) showed how the construction of the SerrePonçon dam in the headwaters of the River Durance in Provence had entirely changed the ecosystem of this steep gradient river. The barrage was completed in 1959. The pre-impoundment river experienced enormous seasonal variations, with very low flows in the summer droughts to devastating floods (6,000 m3 /s−1 ). The Serre-Ponçon dam—with its 1270 million m3 of storage—now completely regulates the flow, but it is only one of a series of dams on the Durance. The

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of north and north-west Europe and even there they raise three key questions. First, which nature should it be? Many states of nature can be created—prehistoric, pre-industrial, or even pre-glacial (Hull and Robertson 2000). Second, can universal indicators be devised that are suitable for most aspects of river restoration in different environments, cultures, and socio-economic constraints? Third, can the rules, values, objectives, and procedures developed for the temperate rivers of northwest Europe be sensibly applied to the seasonally controlled humid and ephemeral rivers of the European Mediterranean?

flow is now more regular, with a diminished snowmelt flood in spring. Several new biotopes have appeared in the channel and floodplain complex and a significant increase in riparian vegetation, especially reed beds, has resulted in an increase in animal and plant diversity in the river corridor. The modified flow regime has also created a new corridor for the spread of plant and animal species. The changes in biodiversity highlight the need for good management and careful attention to the impact of changes on river ecosystems. This example also highlights a common problem throughout the Mediterranean region—the tension between what is desirable for mountain catchment areas and the resource needs of the littoral plains where economic activities and population are concentrated (Chapter 21).

Concluding Comments: Problems and Prospects

The Water Framework Directive

In the seasonal rivers of the Mediterranean there are quite different management priorities from those of temperate Europe. These include:

In December 2000 the European Union published the Water Framework Directive (Directive 2000/60/CE). This is a legal requirement on all member states to create an information base and an implementation plan so that all inland and coastal waters will reach ‘good status’ by 2015. They will do this by establishing a river basin district structure within which demanding environmental targets will be set, including ecological measures, for surface waters. Monitoring networks for water quantity and quality are patchy across the Mediterranean region and this represents a very significant management problem. The paradigm is now established and we briefly consider here whether it provides an appropriate model for Mediterranean rivers, either those already within the Union, directly or by association, or for those outside it. This Directive reflects the emergence of three important paradigms of river management, worldwide. First, rather than seeking to continue the modification of rivers that has taken place throughout the last two millennia, river managers should seek to restore rivers as close as possible to ‘natural’ conditions (Brookes and Shields 2001). Second, to do this, the river must be seen as a product of its catchment and integrated catchment management (ICM) must be adopted. Consequently changes may be required in the catchment in order to restore the river. Third, this is part of an overall paradigm towards restoring nature that has come to the forefront of the environmental movement (Gobster and Hull 2000). The Water Framework Directive carries with it extensive documentation, including specific guidelines about how to proceed with the implementation. These are more easily implemented for the temperate river basins

r Many rivers have high sediment yields, so the focus

of interest in ICM is the erosion and transport of particulate matter in the catchment. Interest therefore has to focus as much on hillslope as channel process and more on semi-natural vegetation and agricultural land use than on recreation and forestry. r The seasonality of flow and the existence of large flood events shift the thrust of the ecological argument as much to the channel beds and the flood corridors as well as to the water quality issues. r Flooding is a high priority and consequently so is flood protection. The flashiness of Mediterranean rivers is well documented and deserves a high profile in restoration works. Insensitive elimination of water control structures in an overenthusiastic response to catchment management could destroy the work of centuries with disastrous outcomes. r In several important examples in the Mediterranean, irrigation requirements are coupled to inter-basin transfers by the question ‘What is the catchment?’ Different qualities of water may become mixed and in the recent lively debate in Spain over the proposed water transfer from the Ebro to the Segura catchments, it has sometimes been forgotten that the Ebro waters are naturally saline at the proposed take-off point as a result of the contribution from the highly saline marls and evaporite rocks of the middle and lower part of the basin just east of Zaragoza. The Segura River in Murcia illustrates all these issues. It has been transformed by human activity for over two

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millennia, from changes of land use in the Bronze Age to the recent water quality implications of the fruitprocessing industry in the towns along its course, that have resulted in a seriously polluted river. Many of the structural changes to its course were devised as a compromise between evacuating flood waters as fast as possible and encouraging aquifer recharge as much as possible. Significant areas of the catchment are severely eroded and there already exists a management plan to reduce erosion and its associated high runoff rates and sediment yield. The flooding problem is an important issue and dam construction is a major element in runoff control and flood protection. The Water Framework Directive represents a crucial shift in thinking towards the principles of integrated catchment management. However, the practice of integrated catchment management is broader than the practice of integrated water management which has been a keystone for many years in the control of Mediterranean rivers. Guidelines cannot be too prescriptive, but must take account of the diversity of both environment and culture in Europe. At the same time they must be rigid enough to ensure compliance. The implementation of the Water Framework Directive is a very costly and demanding activity to impose on those catchments that are thinly populated or in economically marginal areas. In closing this chapter we reiterate the point that it is impossible to understand the hydrology of Mediterranean catchments without appreciating the huge impact of human activities. On the other hand, nature is a strong force that cannot be ignored. Incursions into areas with a high flood risk, careless disregard for modifications of land use and exploitative use of groundwater all come at a high cost. Whilst it can be argued that the integrated approach of the Water Framework Directive in the dryland catchments of southern Europe would be inappropriate (Thornes and Rowntree 2006), it might yet yield important lessons for the integrated management of river catchments across the entire Mediterranean region.

Acknowledgements The authors thank the external reviewer who provided valuable comments on an earlier draft of this chapter. Nick Scarle in the Cartographic Unit at The University of Manchester expertly redrafted the figures. We thank Ian Foster for providing photographs of Israel and Tunisia.

References Beven, K. (2002), Run-off generation in semi-arid areas, in L. J. Bull and M. J. Kirkby (eds.), Dryland Rivers: Hydrology

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and Geomorphology of Semi-Arid Channels. John Wiley & Sons, Chichester, 57–105. Bochet, E., Poesen, J., and Rubio, J. L. (2006), Runoff and soil loss under individual plants of a semi-arid Mediterranean shrubland: influence of plant morphology and rainfall intensity. Earth Surface Processes and Landforms, 31/5: 536–50. Borel, L. (1994), Influence des aménagements sur l’évolution des Milieux Durnciens—dynamique des peuplements vegetaux et animaux, in J. Riser (ed.), Aménagements et gestion des grandes rivières mediterranéennes. Étude Vauclusiennes 5: 15–19. Brookes, A. and Shields, F. D. (2001), River Restoration. John Wiley & Sons, Chichester. Bull, L. J., Kirkby, M. J., Shannon, J., and Hooke, J. M. (2000), The impact of storms on floods in ephemeral channels in south-east Spain. Catena 38/3: 191–210. Butcher, G. C. and Thornes, J. B. (1978), Spatial variability in runoff processes in an ephemeral channel. Zeitschrift für Geomorphologie Suppl. 29: 83–92. Carson, M. A. and Kirkby, M. J. (1972), Hillslope Form and Process. Cambridge University Press, Cambridge. Chirino, E., Bonet, A., Bellot, J., and Sanchez, J. R. (2006), Effects of 30 year old Allepo pine plantations on runoff, soil erosion and plant diversity in a semi-arid landscape in south-eastern Spain. Catena 31/1: 19–30. Christopoulos, G. (1998), Late Holocene river behaviour in the lower Alfios basin, Western Greece. Ph.D. Thesis, University of Leeds. Conesa-Garcia, C. and Alonso-Sarria, F. (1997), Stochastic matrices applied to the probabilistic analysis of runoff events in a semi-arid stream. Hydrological Processes 11: 297–310. Cudennec, C., Leduc, C., and Koutsoyiannis, D. (2007), Dryland hydrology in Mediterranean regions—a review. Hydrological Sciences Journal 52: 1077–87. Dedkov, A. P. and Moszherin, V. T. (1992), Erosion and sediment yield in mountain areas of the world, in D. E. Walling, T. R. Davies, and B. Hasholt (eds.), Erosion, Debris Flows and Environment in Mountain Regions of the World. IAHS Publication 209: 29–36. Dunne, T. and Black R. D. (1970), Partial area contributions to storm runoff-producing zones in a small New England watershed. Water Resources Research 6: 1296–311. Eagleson, P. S. (1978), Climate, soil and vegetation. 6. Dynamics of the annual water balance. Water Resources Research 14/5: 749–64. Elwell, H. A. and Stocking, M. A. (1976), Vegetation cover to estimate soil erosion hazard in Rhodesia. Geoderma 15: 61–70. Evenari, M., Shanan, L., and Tadmor, N. (1982), The Negev: The Challenge of a Desert. Harvard University Press, Cambridge, Mass. Francis, C. F. and Thornes, J. B. (1990), Runoff hydrographs from the Mediterranean vegetation cover types, in: J. B. Thornes (ed.), Vegetation and Erosion, John Wiley and Sons, Chichester, 313–84. (1994), Matorral: erosion and reclamation, in J. Albaladejo, M. A. Stocking, and E. Diaz (eds.), Soil Degradation and Rehabilitation in Mediterranean Environmental Conditions. Consejo Superior de Investigaciónes Científicas, Madrid, 87–116. Garcia, C., Laronne, J. B., and Sala, M. (2000), Continuous monitoring of bedload flux in a mountain gravel-bed river. Geomorphology 34: 23–31.

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Garcia-Ortiz, E. (2006), Efecto de la structura de la copa en la partición de lluvia de tres species arbustivas en clima semiarido. Doctoral thesis, University of Almeria. Gil Olcina, A. (1971), El Campo de Lorca: estudio de geografía agraria. CSIC, Valencia. Gobster, P. H. and. Hull, R. B (eds.) (2000), Restoring Nature: Perspectives from the Social Sciences and Humanities. Island, Washington. Gómez-Espin, J. M. (2004), Approvechamiente Integral del Agua en la Rambla de Nogalte (Puerta Lumbreras-Murcia), University of Murcia, Murcia. Grenon, M. and Batisse, M. (eds.) (1989), Futures for the Mediterranean Basin: The Blue Plan. Oxford University Press, Oxford. Grove, A. T. and Rackham, O. (2003), The Nature of Mediterranean Europe: An Ecological History. Yale University Press, New Haven. Hull, R. B. and Robertson, D. P. (2000), The language of nature matters: we need more Public Ecology, In P. H. Gobster and R. B. Hull (eds.), Restoring Nature: Perspectives from the Social Sciences and Humanities. Island, Washington. Imeson, A. C., Verstraten, E. J., Mulligan, J. M., and Sevink, J. (1992), The effects of fire and water repellency on infiltration and runoff under Mediterranean type forest. Catena 19: 345–61. Inbar, M. (1992), Rates of fluvial erosion in basins with a Mediterranean type climate. Catena 19: 393–409. Laronne, J. B. and Reid, I. (1993), Very high rates of bedload sediment transport by ephemeral desert rivers. Nature 366: 148–50. Kapsimalis, V., Poulos, S. E., Karageorgis, A. P., Pavlakis, P., and Collins, M. (2005), Recent evolution of a Mediterranean deltaic coastal zone: human impacts on the Inner Thermaikos Gulf, NW Aegean Sea. Journal of the Geological Society, London 162: 879–908. Kirkby, M. J. (1969), Infiltration, throughflow and overland flow, In R. J. Chorley (ed.), Water, Earth and Man. Methuen, London, 215–28. Bracken, L., and Reaney, S. (2002), The influence of land use, soils and topography on the delivery of hillslope runoff to channels in SE Spain. Earth Surface Processes and Landforms 27: 1459–73. Lahlou, A. (1988), The silting of Moroccan dams, in M. P. Bordas and D. E. Walling (eds.), Sediment Budgets. IAHS Publication 174: 71–7. Lange, J., Leibundgut, C., Greenbaum, N., and Schick, A. P. (1999), A non-calibrated rainfall-runoff model for large arid catchments. Water Resources Research 35: 2161–72. Lewin, J., Macklin, M. G., and Woodward, J. C. (eds.) (1995), Mediterranean Quaternary River Environments. Balkema, Rotterdam. Macklin, M.G., Lewin, J., and Woodward, J. C. (1995), Quaternary fluvial systems in the Mediterranean basin, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Balkema, Rotterdam, 1–25. McNeill, J. R. (1992), The Mountains of the Mediterranean World. Cambridge University Press, Cambridge. Mavromati, E. and Chryssaidis, L. (2007), Aqueducts in the Hellenic area during the Roman Period. Water Science and Technology: Water Supply 7: 139–45. Nicholas, A. P., Woodward, J. C., Christopoulos, G., and Macklin, M. G. (1999), Modelling and monitoring the impact of dam construction and gravel extraction on rates of bank

erosion in the Alfios River, Peloponnese, western Greece, in A. G. Brown and T. A. Quine (eds.) Fluvial Processes and Environmental Change. John Wiley & Sons, Chichester, 117–37. Obando, J. (2002), The impact of land abandonment on regeneration of semi-natural vegetation. A case study from the Guadalentin, in J. A. Geeson, C. J. Brandt, and J. B. Thornes, Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 269–77. Pocklington, R. (1986), Acequias árabes y pre-árabes en Murcia y Lorca: apportación toponímica a la historia del regadío, in X Colloqui General de la Societat d’Onomàstica, 1985. University of Valencia, 462–73. Probst, J. L., and Amiotte Suchet, P. (1992), Fluvial suspended sediment transport and Mechanical erosion in the Maghreb (North Africa). Hydrological Sciences Journal 37: 621–37. Reid, I. and Laronne, J. B. (1995), Bedload sediment transport rates in an ephemeral stream and a comparison with seasonal and perennial counterparts. Water Resources Research 31: 773–81. and Powell, D. M. (1998), Flash-flood and bedload dynamics of desert gravel-bed streams. Hydrological Processes 12: 543–57. Rojo Serrano, L., Garcia Robredo, F., Martinez Artero, J. A., and Martinez Ruiz, A. (2002), Management Plan to combat desertification in the Guadalentin River Basin, in N. A. Geeson, C. J. Brandt, and J. B. Thornes (eds.), Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 303–19. Rovira, A., Batalla, R. J., and Sala, M. (2005), Fluvial sediment budget of a Mediterranean river: the lower Tordera (Catalan Coastal Ranges, NE Spain). Catena 60: 19–42. Shannon, J., Richardson, R., and Thornes, J. B. (2002), Modelling event-based fluxes in ephemeral streams in L. J. Bull and M. J. Kirkby (eds.), Dryland Rivers: Hydrology and Geomorphology of Semi-Arid Channels. John Wiley and Sons, Chichester, 129–72. Struglia, M. V., Mariotti, A., and Filograsso, A. (2004), River discharge into the Mediterranean Sea: climatology and aspects of the observed variability. Journal of Climate 17: 4740–51. Syvitski, J. P. M., Vörösmarty, C. J., Kettner, A. J., and Green, P. (2005), Impact of humans on the flux of terrestrial sediment to the global coastal ocean. Science 308: 376–80. Thornes, J. B. (1977), Channel changes in ephemeral streams: observations, problems and models, in K. J. Gregory (ed.), River Channel Changes. John Wiley & Sons, Chichester. (2007), Modelling soil erosion by grazing: recent developments and new approaches. Australian Geographical Research 45: 13–26. and Rowntree, K. (2006), Integrated catchment management in semi-arid environments in the context of the European Water Framework Directive. Land Degradation and Development 17: 255–264. Victoria Jumilla, F. and Vicente Lopez, E. (1986), La contaminación de las aguas en la region de Murcia. Ministerio de Obras Publicas, Murcia. Vita-Finzi, C. (1978), Archaeological Sites in their Setting. Thames & Hudson, London. Wainwright, J. (1996), Infiltration, runoff and erosion characteristics of agricultural land in extreme storm events, S. E. France. Catena 26: 27–47.

Hydrology, River Regimes, and Sediment Yield and Thornes, J. B. (2003), Environmental Issues in the Mediterranean: Processes and Perspectives from the Past and Present. Routledge, New York. Walling, D. E. (1984), The sediment yields of African rivers, in D. E. Walling, S. S. D. Foster, and P. Wurzel (eds.), Challenges in African Hydrology and Water Resources. IAHS Publication 144: 265–83. Woodward, J. C. (1995), Patterns of erosion and suspended sediment yield in Mediterranean river basins, in I. D. L. Foster, A. M. Gurnell, and B. W. Webb (eds.), Sediment and Water

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Quality in River Catchments. John Wiley & Sons, Chichester, 365–89. and Foster, I. D. L. (1997), Erosion and suspended sediment transfer in river catchments: environmental controls, processes and problems. Geography 82/4: 353–76. Macklin, G., Krom, M. D., and Williams, M. A. J. (2007), The Nile: evolution, Quaternary river environments and material fluxes, in A. Gupta (ed.), Large Rivers: Geomorphology and Management. John Wiley & Sons, Chichester, 261–92.

This chapter should be cited as follows Thornes, J. B., López-Bermúdez, F., and Woodward, J. C. (2009), Hydrology, river regimes, and sediment yield, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 229–253.

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9

Lakes, Wetlands, and Holocene Environmental Change Neil Roberts and Jane Reed

Introduction The Mediterranean regions of the world are defined on the basis of their climate, with a distinct hot, dry summer season and a warm, wet winter (Grove and Rackham 2001; Chapter 3). Spring and autumn seasons are less well defined but often contribute significantly to annual precipitation. Strictly defined in this way, the Mediterranean region is confined to parts of Italy, Greece, southern France, the south and east of Spain (non-Atlantic climate), the Maghreb and Cyrenaica in North Africa, and narrow coastal strips running through the Balkans, southern and western Turkey, and the Levant (Syria, Lebanon, and Israel-Palestine) (Figure 9.1a). Outside these areas, climate becomes humid temperate (western Europe, Black Sea), arid (Sahara, northern Arabia), or continental (interior areas of the Balkans, Turkey and Iberia, the Zagros mountains of Iran/Iraq). Even within the strict definition are found subalpine mountain zones, so it is a difficult study region to demarcate absolutely. In a similar vein to the volume by Zolitschka et al. (2000), this chapter extends the scope to important wetlands in some neighbouring regions, and deals effectively with the circum-Mediterranean. Thus, we include lakes Ohrid and Dojran in the Balkans, wetlands of the continental interior of Turkey, north-western Iran and the Caucasus (e.g. Lakes Van, Urmia, and Sevan), the climatically dry Jordan rift valley which includes the Dead Sea, and the subalpine northern Italian lakes such as Como and Maggiore. The Mediterranean basin is geologically complex and has its origin in the progressive closure of the Sea of

Tethys during the Tertiary (Laubscher and Bernoulli 1977). Plate convergence between Africa and Eurasia led to a major phase of orogenesis and the creation of fold mountains including the Atlas, Sierra Nevada, Alps, Apennines, and Taurus, and to plateau uplift in Iberia and Anatolia (Chapter 1). These mountain ranges are commonly dominated by massively deformed Mesozoic limestones that now form karst landscapes (e.g. Dinaric Alps; Ager 1980; Chapter 10). Tectonic movement also led to extensive late Cenozoic volcanism, notably in southern and central Italy, the Hellenic arc, Anatolia, and around the Jordan rift (Chapter 15). The Alpine orogeny created not only mountain ranges, but also numerous intervening sedimentary basins, many of which are now occupied by lakes that can be large and sometimes deep (Chapter 4). This geological background is crucial to understanding lake genesis, and has led to important local and regional diversity in topography, rock structure and composition (and hence groundwater hydrology and surface water chemistry), and climatic conditions. Our definition of wetlands includes lakes, marshes, and coastal lagoons, with a focus on natural lakes. The high variability in Mediterranean wetland types has implications both in terms of strategies for conservation and for interpretation of their palaeolimnological records. Different lakes exhibit different thresholds of response to a given forcing function, and there is an unusually great need to understand each lake individually before moving on to attempt any regional syntheses of, for example, lakes as signals of regional climate change.

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17 18

13

27 15 16

24 14

23

25

28

29

30 31

26

32

22

Fig. 9.1. Maps showing: (a) Mediterranean type climates; (b) the location of the largest lakes in the circum-Mediterranean region >200 km2 in area (see Table 9.1 for details); (c) exemplar lake types (see Table 9.2 for details); (d) location of selected key Holocene palaeolimnological sites (see Table 9.3 for details).

Taxonomy and Distribution of Mediterranean Lakes and Wetlands The origins, geomorphology, hydro-chemistry, and biology of any lake, lagoon, or marsh are intimately interlinked with each other and with their geologic and climatic setting (Timms 1992; Wetzel 2001). Mirroring the topographic and geological complexity of the region, Mediterranean wetlands have varied origins when compared to many other ‘lake districts’ such as those of northern Europe and North America which are dominated more uniformly by lakes that are the legacy of Pleistocene glaciation. The location, origins, and main morphometric parameters of a range of Mediterranean lakes are described in Tables 9.1 and 9.2, along with some of the main human impacts upon them. They include many of the designated Ramsar sites (Ramsar 2004) and key sites monitored by the International Lake Environment Committee (ILEC; ). These lakes have been selected to represent the overall variety of different types in the circum-Mediterranean region, along with those that have been the subject of intensive palaeoenvironmental research (Table 9.3) or those that

are large and/or have high economic or ecological status (Table 9.4).

Genesis and Morphology Tectonic Many Mediterranean lakes and wetlands are tectonic in origin. One of the most remarkable is Lake Ohrid on the border between Albania and Macedonia (Former Yugoslav Republic of Macedonia: FYROM). This is Europe’s best example of a steep-sided, quadrangular graben, formed by parallel faulting and subsidence since the Late Miocene (Stankovi´c 1960). On a par with Baikal, Tanganyika, and Malawi, it is ranked amongst the most ancient lakes in the world, having been formed at least 4 million years ago. Other grabens include the nearby but smaller lakes of Prespa and Dojran, or the Greek lakes of Trichonis and Amvrakia. Most other Mediterranean ‘great lakes’ also owe their origin—at least in part—to major long-term tectonic faulting. There are good examples of fault-bounded lakes in Turkey including Sapanca and Iznik (north-west), Burdur, E˘girdir, and Bey¸sehir (south-west), and Hazar (south-east). The Jordan is the most famous of all rift valleys in the Mediterranean region and this contains the

TABLE 9.1 Individual characteristics of all permanent natural lakes in the circum-Mediterranean region >200 km2 in area, excluding coastal lagoons Name

Country

Genesis

Lat N Long E Altitude masl

Lake area km2

Catchment area km2

Max depth m

Salinity g/l

Notes

Sources

Ramsar, Holocene lake sediment record Sediments varved, Late Pleistocene–Holocene sediment cores Ramsar, switch from oligo–to meso-trophic, 19 m lake-level fall since 1933 Falling water levels, southern basin now isolated and used for salt pans, Pleistocene Lisan beds exposed around lake Significant commercial fisheries Former important crayfish production, now mostly lost Ramsar, largest bird colony in Europe, incl. pelicans Lake tourism, partial recovery from pollution World Heritage Site, oligotrophic

Schweizer 1975; Kelts and Shahrabi 1986 Kempe 1977; Degens et al. 1984, Wick et al. 2003

1. Urmia (Rezaiyeh) 2. Van

Iran

Tectonic

37 30 45 30

1280

4750

51000

13

95

Turkey

Lava-dammed

38 30 43 00

1648

3574

12522

451

21

3. Sevan

Armenia

Tectonic

40 20 45 20

1905

1360

3390

86

20◦ C) spring deposits (see Glover and Robertson 2003). Both types are common in the Mediterranean region reflecting the widespread presence of carbonate host rocks with welldeveloped karstic drainage—and the particular circumstances where active tectonic settings allow such groundwaters to be heated by geothermal energy. The Mediterranean region contains some of the best-known examples of Pliocene and Quaternary age and modern tufa deposits because the physical geography of many localities in the region yields optimum conditions for tufa formation (Figure 10.16). These processes take place at a range of scales. Large tufas may form in tectonically controlled depositional settings that allow extended periods of tufa deposition (Glover and Robertson 2003). These may incorporate many facies types with rich fossil assemblages and, because they can be dated by uranium-series methods, they may represent very significant archives of environmental change. The coastal city of Antalya and its hinterland in south-east Turkey (Figure 10.1) sits on one of the largest (c.600 km2 ) and thickest (up to 250 m) tufa deposits in the world (Glover and Robertson 2003) (Figure 10.16a). These workers argue that this large tufa complex was deposited in a large tectonically controlled basin between c.2 and 1.5 million years BP within a series of small lakes when the karstic springs were much more active than at present. The Antalya tufa is composed of pure low-Mg calcite and this is typical of cool-water tufas in the Mediterranean region. Tectonic uplift of the tufa mass created an upper 300 m terrace and glacio-eustatic sea-level change in the Early to Middle Pleistocene created a lower terrace (c.100– 200 m a.s.l.). Subsequent fluvial activity has cut deep gorges in the tufa and the coastal sections have been eroded by wave action (Glover and Robertson 2003). Thus, the present landscape is the product of complex interactions between a wide range of karstic and nonkarstic geomorphological processes and materials over the last two million years or so. In many places the surface of the tufa is crusted and displays complex secondary karst forms following dissolution and recementation of the tufa deposits (Glover and Robertson 2003). In many areas deep terra rossa type soils have developed. The deposition of secondary carbonates may form cemented alluvial gravels. These are widespread in Mediterranean karst environments and the

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(a)

(b)

Fig. 10.14. Two limestone gorges at different stages of development in the Mediterranean region. Both are the product of fluvial incision and subaerial weathering of the bedrock cliffs. (a) The Moracha Canyon in Montenegro and (b) the Vikos Gorge in north-west Greece (photos: Jamie Woodward).

best-developed records are typically Pleistocene in age (Vita-Finzi 1969; Hamlin et al. 2000; Macklin et al. 2002; Woodward et al. 2008). This can help preserve the alluvial record and has allowed the development of long-term records of Pleistocene river behaviour through the application of uranium-series dating methods (Chapter 11). Figure 10.16b shows a section in over 20 m of cemented Pleistocene alluvial gravels in one of the wider reaches of the Moracha Canyon in Montenegro to the north of Podgorica (Figure 10.1). Ambient temperature deposited carbonates in streams have led to the formation of tufa barrages, cascades, or sheets. These can be impressive and deposition can take place above growing barriers, as in the case of the River Krka in Croatia where there is an alluvial plain above the 20-m high Topolje Falls (Jennings 1985), whilst the

River Korana is divided into sixteen separate lakes along a 12-km gorge at Plitvice (Figure 10.1). Thermogene travertines are developed, again spectacularly, in Turkey at Pamukkale, at the Roman city of Hierapolis where 35–39◦ C waters emerge at the surface to form a very distinctive topography (Figure 10.16c). The thick travertines at Bagni de Tivoli near Rome have been quarried for over 2,000 years. Travertine, along with marble, was widely used as a building material in Classical Rome because it is strong yet relatively light due to its open texture. These carbonate rocks represent the foundation and essence of classical architecture. Pentecost (1995) provides a useful catalogue of both ambient-temperature and thermogene deposits in Europe and Turkey, whilst sites in Mediterranean North Africa are discussed by Ford and Pedley (1996).

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Fig. 10.15. Relict karst pinnacles in the White Desert of Egypt (photo: Tony Waltham).

Karst Systems and Environmental Change Given the marked environmental changes that took place in the Mediterranean region during the Pleistocene and Holocene (Chapters 2, 4, 9, 11, and 12), it is clear that karst morphogenesis has been considerably affected by factors other than the autogenic effects so far discussed as, for example, in the cave evolution model presented in Figure 10.7. In broad terms, it is possible to identify four major environmental changes that have impacted upon the karst systems of the region. First, the elevation and exhumation of limestone surfaces, river basin evolution, and karst development have proceeded alongside continuing tectonic activity in the Mesozoic and Cenozoic eras (e.g. Harvey and Wells 1987; Chapter 1) and, initially, under a climate warmer than present (Chapter 4). In southern France, bauxite deposits (named from the cave-ridden spur of Les Baux) were created by karst weathering even before submergence in a Late Cretaceous sea (Ager et al. 1980). Second, the availability of relief has varied, both in response to eustatic Quaternary sea level change (of the

order of 120 m) and earlier in response to the remarkable Messinian (Late Miocene) drying of the Mediterranean Sea basin (Hsü 1972; Adams et al. 1977), when regional base levels fell by more than 1,000 m (Chapter 1). Limestone areas that are now coastal, or near to sea level, were elevated above the littoral zone during the Messinian Salinity Crisis and, repeatedly, during Pleistocene cold-stage low sea-level stands. This means that karst forms are now ‘drowned’ deep in the vadose zone. Recently, Bakalowicz et al. (2003) have considered the development of the deep karst aquifers in the Mediterranean coastal zone that formed during the Late Miocene Messinian Salinity Crisis around 5.5 Ma. The sea-level fall during this period led to the development of large canyons in the lower reaches of the major river systems. This base level fall also led to the formation of very deep karst systems that lie well below present sea levels. The residence time of the groundwaters stored in these deep coastal karsts is not known. These deep karst aquifers have very long vertical conduits and are unique to the Mediterranean region (Mocochain et al. 2006; Bakalowicz et al. 2008). From a resource management perspective, it is important to establish

(a) (b)

(c)

Karst Geomorphology

309

Fig. 10.17. Karst terrain in north-east Majorca showing bare limestone slopes and thick terra rossa soils on the valley floor (photo: Jamie Woodward).

their water resource potential and their role in the contemporary hydrological cycle. Bakalowicz et al. (2003) have discussed the hydrogeological significance of these deep aquifer systems and the difficulties involved in monitoring submarine karst springs and the dynamics of freshwater and sea water exchange in the deep phreatic zone. In Dalmatia these springs are known as vrulje and they have been reported from around the Mediterranean coastal karst. Fleury et al. (2008) monitored the behaviour of two submarine outlets from a karstic system in southern Spain during the 1999–2000 hydrological year. One of the conduits was highly responsive to rapid infiltration across the terrestrial catchment during highintensity rainfall events. Fig. 10.16. (opposite) Three landscape features in the Mediterranean produced by the precipitation of secondary carbonates. (a) The coastal cliffs in tufa deposits at Anatalya on the southern coast of Turkey. (b) A 20-m exposure in cemented alluvial gravels of Pleistocene age in the lower reaches of the Moracha River in Montenegro. (c) The dramatic pools and slopes at Pamukkale in south-west Turkey. All locations are shown on Figure 10.1. (Photos: (a) and (c), Tony Waltham; (b), Jamie Woodward).

Third, the effects of Pleistocene climate fluctuations have been superimposed upon the broad framework of tectonic activity across the region to impose, in very general terms, a series of wet/dry and cold/warm phases on karstic systems (Chapters 2, 4, 11, and 12). These shifts in Quaternary climate altered the balance between physical and chemical weathering in limestone areas, with more intense physical weathering (and in places more extensive cirque and valley glaciation) of exposed upland limestones during cold episodes (Chapter 12). Thus cliff weathering in many limestone gorges produced Pleistocene talus slopes and boulder-bed streams in excess of those generated by cliff collapse and rockfall today. Lithological and mineralogical analysis of Pleistocene alluvial materials suggests much greater incorporation of physical weathering products in cold phase deposits than later (Lewin et al. 1991; Woodward et al. 1992). The Quaternary also involved periods of higher and lower water tables, so that valley networks and river reaches, and cave systems that are now dry were being actively developed by flowing water. Groundwater bodies beneath some of the desert areas of North Africa were

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TABLE 10.3. Characteristics of active karst settings and passive karst settings for rockshelter and cave entrance environments in limestone terrains Active Karst Setting r Linked to an internal cavern or conduit system r Dripping vadose waters r Seasonal water flows and ponding r Range of hydrological pathways r Precipitation of calcite and other minerals r Inwashing of fine sediments via conduits in the host bedrock r Development of vegetation within the site r Mineralization of macroscopic plant remains r Strong chemical diagenesis and mineral alteration r Humidity may encourage host rock breakdown by frost action r Evidence of erosion and sediment removal by invasive karst waters Passive Karst Setting r No significant links with an internal conduit system r Dry site without flowing or dripping water r Limited inwashing of sediments via karstic cavities r Highly localized or no chemical precipitation r Limited vegetation growth in the site r Desiccation of macroscopic plant remains r Limited chemical diagenesis and mineral alteration r Import of fine sediments through the shelter opening may be dominant r Limited host rock weathering by solution r Subaerial processes are dominant Note: See Fig. 10.8. Sites may shift between these two end-members in response to climate change. Source: Modified from Woodward and Goldberg (2001).

developed under more humid Pleistocene conditions, so that they now constitute a fragile, fossil resource that is not effectively being replenished. Fourth, the removal of a soil and vegetation cover, particularly following human activity during the course of the Holocene (Chapters 6 and 9), has transformed karst systems. Thus deep weathering forms developed under a soil cover (known as rundkarren and generally of rounded form with residual core stones) have become exposed on the surface to form small-scale landscapes of pinnacles, blocks, and chasms as shown in the karst terrain of central Majorca in Figure 10.17. These terrains may be modified by the development of sharp-edged flutes (rillenkarren) and channels (rinnenkarren) produced subsequently by the effects of subaerial solution on bare limestone surfaces. Thus some karst features may be composite forms that owe their origin to a combination of different climatic episodes (and associated process regimes involving exhumation and/or burial by Quaternary sediments) just as cave systems may develop under different phases of water flow and chemical balance (Figures 10.7 and 10.8). Complex surface topography, with a combination of relict and active forms, may have an exposed local relief of

the order of 10 m, as in the ‘roches ruiniformes’ of the French Causses or the ‘cuidad encantada’ of the Sierra delibar in southern Spain.

Archives of Environmental Change Sediments in Limestone Rockshelters and Cave-mouth Environments Caves and rockshelters are found wherever hard limestones are present in the Mediterranean basin. Most deep caves owe their formation to karstic processes to a greater or lesser extent, while some shallow rockshelters may result from non-karstic processes such as fluvial erosion and undercutting, or physical weathering of a gorge wall, for example. In the same way the sediments that accumulate in cave-mouth and rockshelter settings will commonly represent some combination of karstic and non-karstic processes—although, in practice, the division is not always clear cut. Woodward and Goldberg (2001) have used the geomorphological and hydrological characteristics of cave-mouth and rockshelter environments to describe active karst settings and passive karst settings (Table 10.3), because the local context is a very important control on the nature of the accumulating sediment. Caves and rockshelters in the Mediterranean have provided shelter for both animals and humans throughout much of the Quaternary and they represent a major archaeological resource. Most of the work on the clastic sedimentary records in rockshelter and cave-mouth environments has been carried out as part of archaeological excavations of Palaeolithic and Mesolithic records (ibid.). Figure 10.18 illustrates the wide range of processes that can transport fine-grained sediments to rockshelter and cave mouth environments in karst settings (Woodward and Bailey 2000). It is also important to appreciate, however, that these processes

Fig. 10.18. (opposite) Rockshelter and cave entrance environments can form important sediment sinks and they represent a major archaeological resource. (a) The opening of Asprochaliko rockshelter in the Louros Valley of north-west Greece. This site contains evidence of Middle and Upper Palaeolithic occupation and was excavated by Eric Higgs and his team in the 1960s (photo: Jamie Woodward). (b) Schematic cross-section of a Mediterranean limestone rockshelter showing the processes that can deliver fine sediments to the site and the natural and archaeological materials that can be used for radiometric dating (based on Woodward and Bailey 2000, and Schwarz and Rink 2001). The allogenic (external) sediment sources are shown in white boxes and the autogenic (internal) ones are shown in the shaded box. Compare to Figure 10.8 and Table 10.3.

(a)

(b)

Detrital component in bedrock

Fine sediments washed through joints, bedding planes and larger conduits.

Physical and chemical weathering of the host limestone bedrock. Debris from tool manufacture and other activities.

Colluvial inputs including fine sediments washed down the gorge wall during storm events, wash from adjacent slopes and mass movement.

Aeolian inputs such as far-travelled dust, tephra and local silts and sands.

Older bedrock profiles

Stalactites ‘Soda straw’ stalactites

Human imports of fine sediment on wet feet and animal carcasses. Raw materials for stone tool manufacture.

Windblown detritus

Bones and teeth Flowstone

Fireplace Eroded channel

Weathered roof block

Natural erosion surface

Fallen stalagmite

Sink hole

Old eroded remnant

Direct fluvial input of fine slackwater sediments during large floods. Fine sediment deposition during high lake or sea level stands.

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can also be destructive and may remove parts of the sedimentary record. Gaps in the record may also represent periods of non-deposition. In the case of Franchthi Cave in southern Greece, for example, Farrand (2000) points out that at least half the interval between c.26,850 and 6,000 cal. years BP is not represented by any deposits in the excavated parts of the site. Thus, detailed dating programmes are needed to establish the timeframe and chronological resolution of the sedimentary record (Bailey and Woodward 1997; Schwartz and Rink 2001). Despite the poor stratigraphic resolution of some sites in comparison to other depositional environments, limestone caves and rockshelters commonly contain a range of materials that can be dated (Figure 10.18). These include materials produced by human activity (e.g. charcoal fragments and burnt flints in hearths), karstic processes (speleothems), and allogenic sediments introduced by natural processes (e.g. tephras and aeolian sands). All these materials can be dated using modern scientific methods and these are summarized in Figure 10.18. The use of large-format thin sections to study the microstratigraphy of these deposits has allowed detailed reconstructions of site formation processes (e.g. Goldberg and Bar-Yosef 1998) and recent work on sediment source identification has allowed these records to be placed in their broader geomorphological context (e.g. Woodward et al. 2001).

Speleothems and Quaternary Palaeoclimates Recent work on speleothems from karstic caves in Israel has provided excellent high resolution records of climate change for the eastern Mediterranean region (BarMatthews et al. 1999, 2000, 2003). Soreq Cave is located west of Jerusalem and around 40 km from the Mediterranean Sea (Figure 10.1). It lies close to the modern 500 mm isohyet at about 400 m above sea level. The speleothems from Soreq Cave and from Peqiin Cave in the north of Israel have yielded a continuous and high resolution record of karstic processes and climate change spanning the last 250,000 years (Figure 10.19). The carbonate materials have been dated using high precision uranium-series methods. These records match closely with the marine oxygen isotope record and have demonstrated that many features of the climate record for the last cold stage in the North Atlantic (such as Heinrich Events) also impacted upon the eastern Mediterranean region (Bar-Matthews et al. 1999) (Chapters 2 and 4).

The caverns in these semi-arid karst landscapes are highly sensitive environmental systems that are well coupled to changes in temperature and precipitation— although it is important to point out that a full understanding of the contemporary karst hydrology is needed before the palaeoclimatic record can be fully appreciated (ibid.). A very significant feature of recent work on the speleothem records in Israel has been the exploration of linkages between the high resolution speleothem data and the marine records from the eastern Mediterranean Sea (Bar-Matthews et al. 2003). This research has shown that some periods of sapropel formation were associated with enhanced precipitation across the eastern Mediterranean region. Sapropels are dark, organic-rich horizons that formed periodically on the bottom of the Mediterranean Sea, and are discussed in Chapter 2. Information on Pleistocene and Holocene sea-level change can also be obtained from speleothems where they have developed in caverns in coastal karst systems that are submerged during high sea-level stands (Chapter 13).When these caverns become flooded by sea water, speleothem growth is halted and they become covered with marine biogenic overgrowths (Suri´c et al. 2005). These overgrowths can be dated by radiocarbon to provide geochronologies for the timing of cavern flooding by sea water although the precise timing of inundation may be conditioned by the geomorphological setting and local karst hydrology. At some sites flooding is by fresh groundwater as local water tables rise in accordance with sea-level rise. Thus, cavern shape and elevation, groundwater conduit gradients, and cavern distance from the coast are important factors in site selection. Suri´c et al. (2005) present data from three submerged caves along the Croatian coast where eight speleothems (ranging from 38.5 to 17 m below present mean sea level) were sampled. Speleothem deposition took place in these caves during the last cold stage between c.37,000 and 22,000 years BP (when global sea levels were around 120 m lower than present). These workers have produced a sea-level curve for the Late Pleistocene and Holocene that is in broad agreement with work on the Tyrrhenian coast of Italy and the Mediterranean coast of France. Similarly, Bard et al. (2002) have estimated interglacial (MIS 7) sea levels from speleothems preserved in Argentarola Cave in Italy.

The Conservation of Karst Environments Today, worldwide academic interest in these karst environments must also be coupled with an appreciation of the value and fragility of Mediterranean karst

Karst Geomorphology

313

–9 1

2

3

4

5

–8

d18O % (PDB)

–7

–6

–5

–4

–3 S1

S3

S4

S5

–2 0

20

40

60 80 Age (thousand years)

100

120

140

Fig. 10.19. A high-resolution oxygen isotope record (‰18 O) from speleothems in Soreq Cave, Israel, spanning the last 140,000 years (modified from Bar-Matthews et al. 1999). This record shows how the regional climate has fluctuated between warmer and wetter periods (peaks) to cooler and drier periods (troughs). It also shows that climate change can be abrupt (see Chapters 4, 9, and 11). Marine oxygen isotope stages 1–5 are shown along with sapropels S1, S3, S4, and S5. Sapropel formation in the eastern Mediterranean Sea is discussed in detail in Chapter 2.

landscapes and groundwater resources (Plagnes and Bakalowicz 2002; Bakalowicz et al. 2008). The conservational aspects of limestone and karsts in the region is attracting increasing attention (Gams et al. 1993). This includes threats to groundwater quality by pollution (both point and diffuse sources and waste disposal) and the potentially damaging exploitation of fossil groundwaters, the destruction of valued landscapes and archaeological sites by quarrying, and by urban/recreational development and land use pressure (Figure 10.20). Karstic lakes are often ecologically important sites but are susceptible to eutrophication and habitat degradation (Reed et al. 2008). There are numerous examples of the unwise use of subsurface water systems which

have not been adequately understood, but which have suffered from pollution or resource depletion. Action to protect aquatic environments by the European Commission has involved vulnerability and risk assessment of carbonate (karst) aquifers. The exploitation of limestone terrains is nothing new—as shown by the huge Carrara marble quarries in Italy that were a resource for Renaissance art and architecture—but the scale of recent transformations is unprecedented. Population pressure and the large number of tourists visiting underground cave systems can also threaten cave environments (including their atmosphere, flora, and fauna) without access management. Limestone caves and rockshelters also

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Human activities

Impacts on Mediterranean karsts

Effects Karst terrain

Adjacent non-karst terrain Deforestation

Impoverished ecology

Loss of biota Reduced evapotranspiration

Agriculture Increased runoff and erosion Urbanization and industry

Soil degradation and erosion Sedimentation of caves

Increased sediment discharge Water quality deterioration Acid rain

Quarrying and mining

Landform destruction Waste water discharge

Tourism and recreation

KARST ECOSYSTEM DEGRADATION

Rock and mineral removal

Chemical wastes

Military activities

Soil collapse Wear and tear Water exploitation Reduced allogenic recharge

Spring desiccation

– dams (upstream) – groundwater pumping – dams (downstream)

Water-table lowering

Sea water intrusion

Inundation

Drowned karst systems

Fig. 10.20. Human activities in the Mediterranean region and their potential impact on non-karst and karst terrains (modified from Williams 1993).

TABLE 10.4. Karst sites in the Mediterranean region with World Heritage status Country Croatia France and Spain Slovenia Spain Turkey Montenegro

Karst environment Plitvice Lakes National Park Pyrenees-Mount Perdu Škocjan Caves Altimera Cave Atapuerca Cave Pamukkale Durmitor National Park

Designation Natural Heritage Natural and Cultural Natural Heritage Cultural Heritage Cultural Heritage Natural and Cultural Natural Heritage

Source: Modified from Hamilton-Smith (2004).

preserve valuable records of Quaternary environmental change that include a rich archaeological resource on the history of Palaeolithic and Mesolithic cultures in particular. Limestone caves in the Mediterranean region— including Altimira Cave in Spain and Chauvet in France (Figure 10.1)—contain some of the best-known examples of Upper Palaeolithic cave art and these also require

careful conservation and management. A recent and very positive development has been the designation of World Heritage Site status by UNESCO for several karst and cave sites in the Mediterranean region and these are listed in Table 10.4.

Conclusions Karst terrains are a very significant component of the physical geography of the Mediterranean basin because the effects of both carbonate dissolution and precipitation processes are widespread. It can be argued that carbonate rocks are an important unifying feature of much of the Mediterranean basin. Many of the key mountain ranges such as the Pindus Mountains of Greece, the Apennines of Italy, and the Dinaric Alps of the former Yugoslavia are dominated by uplifted limestone terrains and these landscapes include glaciokarsts, deep

Karst Geomorphology

limestone gorges, large dolines, sinking streams, and extensive cave systems. Limestone bedrock coasts are also a key feature of the Mediterranean environment (Chapter 13). Much of the river water that flows into the Mediterranean Sea has passed through karst terrain, so this environment is an important influence on both river regimes and water quality (Chapter 8). Karst landscapes pose particular engineering problems and ground collapse can be a significant geohazard in many areas. Karst also leads to the formation of distinctive red terra rossa soils (Chapter 6) and vegetation communities and it has influenced the development of agricultural practices and water resource management strategies for thousands of years. In some parts of the Mediterranean region—particularly those in North Africa—the karsts have yet to be fully explored. This is reflected in the much smaller literature on the karsts of several countries including Libya, Algeria, and Tunisia, for example (see entries in Gunn 2004). An understanding of karstic processes is essential for many disciplines and may provide insights in unexpected contexts. For example, recent work at the Early Bronze Age fortified site of Vayia, in southern Greece, has shown that the variable development and orientation of karren features on limestone blocks found within cairns could be used to establish a relative dating framework for architectural remains and to establish that the cairns themselves were formed in antiquity (Tartaron et al. 2006). It is also important to appreciate the significance of carbonate precipitation dynamics in the karst systems of the region. The formation of tufa, travertine, speleothems, and calcrete is also a key feature of Mediterranean karst environments and this has produced a highly distinctive suite of karstic features and landscapes. The Mediterranean basin contains some of the most extensive and spectacular Pliocene- and Quaternary-age tufa deposits in the world (Glover and Robertson 2003) including the distinctive terrains at Pamukkale and Plitvice. Carbonate precipitates often constitute important archives of Quaternary environmental change that can be dated by uranium-series methods. This approach has provided new insights into the Pleistocene glacial record in the Mediterranean and recent work on speleothems in Israel, in particular, has produced very detailed and globally significant records of environmental change for the last 250,000 years (e.g. Bar-Matthews et al. 2003). Mediterranean karst landscapes have been shaped by a range of spatially and temporally variable environmental fluctuations in the recent geological past. In broad terms these reflect the variable impacts of long-term tectonic history and Quaternary climate change and, in the Holocene, the

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increasing intensity of human activity across the basin from Neolithic to modern times.

References Adams, C. G., Benson, R. H., Kidd, R. B., Ryan, W. B. F., and Wright, R. C. (1977), The Messinian salinity crisis and evidence of late Miocene eustatic changes in the world ocean. Nature 269: 383–6. Ager, D. V. (1980), The Geology of Europe. McGraw Hill, London. Albritton, C. C., Brooks, J. E., Issawi, B., and Swedor, A. (1990), Origin of the Qattara Depression, Egypt. Bulletin of the Geological Society of America 102: 952–60. Bailey, G. N. and Woodward, J. C. (1997), The Klithi deposits: sedimentology, stratigraphy and chronology, in G. N. Bailey (ed.), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece, i. Excavation and Intra-site Analysis at Klithi. Cambridge, McDonald Institute for Archaeological Research, 61–94. Turner, C., Woodward, J. C., Macklin, M. G., and Lewin, J. (1997), The Voidomatis basin: an introduction, in G. N. Bailey (ed.), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece, i. Klithi in its Local and Regional Setting. Cambridge, McDonald Institute for Archaeological Research, 321–45. Bakalowicz, M., Fleury, P., Jouvencel, B., Promé, J. J., Becker, P., Carlin, T., Dörfliger, N., Seidel, J. L., and Sergent, P. (2003), Coastal karst aquifers in Mediterranean regions: a methodology for exploring, exploiting and monitoring sub-marine springs. Tecnologia de la Intrusion de Agua de Mar en Acuiferos Costeros: Paises Mediterraneos. IGME, Madrid. El Hakim, M., and El-Hajj, A. (2008), Karst groundwater resources in the countries of eastern Mediterranean: the example of Lebanon. Environmental Geology 54: 597–604. Bard, E., Antonioli, F., and Silenzi, S. (2002), Sea-level during the penultimate interglacial period based on a submerged stalagmite from Argentarola Cave (Italy). Earth and Planetary Science Letters 196: 135–46. Bar-Matthews, M., Ayalon, A., Kaufman, A., and Wasserburg, G. J. (1999), The Eastern Mediterranean palaeoclimate as a reflection of regional events: Soreq cave, Israel. Earth and Planetary Science Letters 166: 85–95. (2000), Timing and hydrological conditions of Sapropel events in the Eastern Mediterranean, as evident from speleothems, Soreq cave, Israel. Chemical Geology 169: 145–56. Ayalon, A., Gilmour, M., Matthews, A., and Hawksesworth, C. (2003), Sea-land oxygen isotopic relationships from planktonic foraminifera and speleothems in the Eastern Mediterranean region and their implication for paleorainfall during interglacial intervals, Geochimica et Cosmochimica Acta, 67: 3181–99. Beckinsale, R. P. and Chorley, R. J. (1991), The History of the Study of Landforms or the Development of Geomorphology 1890 to 1950. New York, Routledge. Biju-Duval, B., Dercourt, V., and Le Pichon, X. (1977), From the Tethys ocean to the Mediterranean sea: a plate tectonic model of the evolution of the western Alpine system, in B. Biju-Duval and L. Montadert (eds.), The Structural History of the Mediterranean Basins. Éditions Technip, Paris, 143–64.

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Bonacci, O., Ljubenkov, I., and Roje-Bonacci, T. (2006), Karst flash floods: an example from the Dinaric karst (Croatia). Natural Hazards and Earth Systems Science 6: 195–203. Brewster, H. (1997), The River Gods of Greece: Myths and Mountain Waters in the Hellenic World. I. B. Tauris, London. Calaforra, J. M. and Pulido-Bosch, A. (2003), Evolution of the gypsum karst of Sorbas (SE Spain). Geomorphology 50: 173–80. Courbon, P. and Chabert, C. (1986), Atlas des Grandes Cavities Mondiales. UIS-FFS. Cviji´c, J. (1893), Das Karstphänomen. Geographische Abhandlungen herausgegeben von A. Penck [Geographical Proceedings. Published by A. Penck] 5/3: 218–329. Dewey, J. F., Pitman III, W. C., Ryan, W. B. F., and Bonnin, J. (1973), Plate tectonics and the evolution of the Alpine system. Bulletin of the Geological Society of America 84: 3137–80. Farrand, W. R. (2000), Depositional History of Franchthi Cave: Sediments, Stratigraphy and Chronology: Fascicle 12. Indiana University Press, Bloomington, Ind. Ferrarese, F., Sauro, U., and Tonello, C. (1998), The Montello Plateau. Zeitschrift für Geomorphologie, Suppl.109: 41–62. Fleury, P., Bakalowicz, M., de Marsily, G., and Cortes, J. M. (2008), Functioning of a coastal karstic system with a submarine outlet, in southern Spain. Hydrogeology Journal 16: 75–85. Ford, D. C. (2004), Karst in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 473–5. and Pedley, H. M. (1996), A review of tufa and travertine deposits of the world. Earth Science Reviews 41: 117–75. and Williams, P. W. (1989), Karst Geomorphology and Hydrology. Unwin Hyman, London. Frumkin, A. (1994), Hydrology and denudation rates of halite karst. Journal of Hydrology 162: 171–89. Gale, S. J. and Hoare, P. G. (1997), The glacial history of the northwest Picos de Europa of northern Spain. Zeitschrift für Geomorphologie, NS 41: 81–96. Gilbertson, D. D., Hoare, P. G., Hunt, C. O., Jenkinson, R. D., Lamble, A. P., O’Toole, C., van der Veen, M., and Yates, G. (1993), Late Holocene environmental change in the Libyan predesert. Journal of Arid Environments 24: 1–19. Gams, I. (1978), The Polje: The problems of its definition. Zeitschrift für Geomorphologie 22: 170–81. Nicod, J., Julian, M., Anthony, E., and Sauro, U. (1993), Environmental change and human impacts on the Mediterranean karsts of France, Italy and the Dinaric region, in P. W. Williams (ed.), Karst Terrains, Environmental Changes, Human Impact. Catena Suppl. 25: 59–98. Gillieson, D. (1996), Caves: Processes, Development and Management. Blackwell, Oxford. Glover, C. and Robertson, A. H. F. (2003), Origin of tufa (coolwater carbonate) and related terraces in the Antalya area, SW Turkey. Geological Journal 38: 329–58. Goldberg, P. and Bar-Yosef, O. (1998), Site formation processes in Kebara and Hayonim Caves and their significance in Levantine prehistoric caves, in T. Akazawa, K. Aoki, and O. Bar-Yosef (eds.), Neandertals and Modern Humans in Western Asia. Plenum, New York, 107–23. Gunn, J. (2004), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London. Hamilton-Smith, E. (2004), World Heritage Sites, in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 777–9. Hamlin, R. H. B., Woodward, J. C., Black, S., and Macklin, M. G. (2000), Sediment fingerprinting as a tool for interpreting long-

term river activity: the Voidomatis basin, NW Greece, in I. D. L. Foster (ed.), Tracers in Geomorphology. John Wiley & Sons, Chichester, 473–501. Harvey, A. M. and Wells, S. G. (1987), Response of Quaternary fluvial systems to differential epeirogenic uplift: Aguas and Feos river systems, southeast Spain. Geology 15: 689–93. Herak, M. and Stringfield, V. T. (eds.) (1972), Karst: Important Karst Regions of the Northern Hemisphere. Elsevier, Amsterdam, 551. Hodge, E. J., Richards, D. A., Smart, P. L., Andreo, B., Hoffmann, D. L., Mattey, D. P., and González-Ramón, A. (2008), Effective precipitation in southern Spain (266 to 46 ka) based on a speleothem stable carbon isotope record. Quaternary Research 69: 447–57. Hsü, K. J. (1972), When the Mediterranean dried up. Scientific American 227: 27–36. Hughes, P. D., Woodward, J. C., Gibbard, P. L., Macklin, M. G., Gilmour, M. A., and Smith, G. R. (2006), The glacial history of the Pindus Mountains, Greece. Journal of Geology 114: 413–34. Jakucs, L. (1977), Morphogenetics of karst regions. Unwin Hyman, London. Jennings, J. N. (1985), Karst Geomorphology. Blackwell, Oxford. Jourde, H., Roesch, A., Guinot, V., and Bailly-Comte, V. (2007), Dynamics and contribution of karst groundwater to surface flow during Mediterranean flood. Environmental Geology 51: 725–30. King, G., Sturdy, D., and Bailey, G. N. (1997), The tectonic background to the Epirus landscape, in G. N. Bailey (ed.), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece. McDonald Institute, Cambridge, ii. 541–58. Klimchouk, A. (2004), Morphometry of Caves, in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 524–6. Lowe, D., Cooper, A., and Sauro, U. (eds.) (1996), Gypsum Karst of the World. International Journal of Speleology (Special Issue) 25: 3–4. Lewin, J., Macklin, M. G., and Woodward, J. C. (1991), Late Quaternary fluvial sedimentation in the Voidomatis Basin, Epirus, northwest Greece. Quaternary Research 35: 103–15. Lowry, D. C. and Jennings, J. N. (1974), The Nullabor karst Australia. Z. Geom. 18: 35–81. Macklin, M. G., Fuller, I. C., Lewin, J., Maas, G. S., Passmore, D. G., Rose, J., Woodward, J. C., Black, S., Hamlin, R. H. B., and Rowan, J. S. (2002), Correlation of Late and Middle Pleistocene fluvial sequences in the Mediterranean and their relationship to climate change. Quaternary Science Reviews 21/14/15: 1633–44. Milanovi´c, P. T. (2000), Geological Engineering in Karst. Belgrade, Zebra. Mocochain, L., Clauzon, G., Bigot, J., and Brunet, P. (2006), Geodynamic evolution of the peri-Mediterranean karst during the Messinian and the Pliocene: evidence from the Ardèche and Rhône Valley systems canyons, Southern France. Sedimentary Geology 188/9: 219–33. Palmer, A. N. (1991), Origin and morphology of limestone caves. Bulletin of the Geological Society of America 103: 1–21. Pena, J. L., Sancho, C., and Lozano, M. V. (2000), Climatic and tectonic significance of Late Pleistocene and Holocene tufa deposits in the Mijares River canyon, eastern Iberian Range, Northeast Spain. Earth Surface Processes and Landforms 25: 1403–17.

Karst Geomorphology Pentecost, A. (1995), The Quaternary travertine deposits of Europe and Asia Minor. Quaternary Science Reviews 14: 1005– 28. Perritaz, L. (2004), Africa, North in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 13–16. and Monbaron, M. (1998), Geomorphological approach to the Aït Abdi Karst Plateau (Central High Atlas, Morocco). Zeitschrift für Geomorphologie 109: 83–104. Plagnes, V. and Bakalowicz, M. (2002), The protection of a karst water resource from the example of the Larzac karst plateau (south of France): a matter of regulations or a matter of process knowledge? Engineering Geology 65: 107–16. Reed, J. M., Leng, M. L., Ryan, S., Black, S., Altinsaçli, S., and Griffiths, H. I. (2008), Recent habitat degradation in karstic Lake Uluabat, western Turkey: A coupled limnological– palaeolimnological approach. Biological Conservation 141: 2765–83. Rossi, C. (2004), Picos de Europa, Spain in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 582–5. Schwarz, H. P. and Rink, W. J. (2001), Dating methods for sediments of caves and rockshelters. Geoarchaeology 16, 355–71. Smart, C. C. and Worthington, S. R. H. (2004), Springs, in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 699–703. Smart, P. L. (1986), Origin and development of glacio-karst closed depressions in the Picos de Europa, Spain. Zeitschrift für Geomorphologie, NS 30; 423–43. Smith, B. J., Warke, P. A., and Moses, C. A. (2000), Limestone weathering in contemporary arid environments: a case study from southern Tunisia. Earth Surface Processes and Landforms 25: 1343–54. Suri´c, M., Juraˇci´c, M., Horvatinˇci´c, N., and Broni´c, I. K. (2005), Late Pleistocene–Holocene sea level rise and the pattern of coastal karst inundation: records from submerged speleothems along the Eastern Adriatic Coast (Croatia). Marine Geology 214: 163–75. Tartaron, T. F., Pullen, D. J., and Noller, J. S. (2006), Rillenkarren at Vayia: geomorphology and a new class of Early Bronze Age fortified settlement in Southern Greece. Antiquity 80: 145–60. Van Houten, F. B. (1980), Latest Jurassic-Early Cretaceous regressive facies, northeast Africa Craton. American Association of Petroleum Geologists Bulletin 64: 857–67. Vita-Finzi, C. (1969), The Mediterranean Valleys: Geological Changes in Historical Times. Cambridge, Cambridge University Press. Waltham, A. C. (1978), The caves and karst of Astraka, Greece. Transacations of the British Cave Research Association 5: 1–12.

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(2001), Pinnacles and barchans in the Egyptian desert. Geology Today 17: 101–4. White, W. B. (1988), Geomorphology and Hydrology of Karst Terrains. Oxford University Press, Oxford. Williams, P. W. (1993), Environmental change and human impact on karst terrains: an introduction, in P. W. Williams (ed.), Karst Terrains, Environmental Changes, Human Impact. Catena Suppl. 25: 1–19. (2004), Karst Evolution, in J. Gunn (ed.), Encyclopedia of Caves and Karst Science. Fitzroy Dearborn, London, 475–8. Windley, B. F. (1984), The Evolving Continents. John Wiley & Sons, Chichester. Woodward, J. C. (1995), Patterns of erosion and suspended sediment yield in Mediterranean river basins, in I. D. L. Foster, A. M. Gurnell, and B. W. Webb (eds.), Sediment and Water Quality in River Catchments. John Wiley and Sons, Chichester, 365–89. (1997), Late Pleistocene rockshelter sedimentation at Megalakkos, in G. N. Bailey (ed.), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece, ii. Klithi in its Local and Regional Setting. McDonald Institute for Archaeological Research, Cambridge, 377–93. and Bailey, G. N. (2000), Terminal Pleistocene sediment sources and geomorphological processes recorded in rockshelter sequences in northwest Greece in I. D. L. Foster (ed.), Tracers in Geomorphology. John Wiley & Sons, Chichester, 521–51. and Goldberg, P. (2001), The sedimentary records in Mediterranean rockshelters and caves: archives of environmental change. Geoarchaeology: An International Journal 16/4: 327–54. Lewin, J., and Macklin, M. G. (1992), Alluvial sediment sources in a glaciated catchment: the Voidomatis basin, northwest Greece. Earth Surface Processes and Landforms 17/3: 205–16. Hamlin, R. H. B., Macklin, M. G., Karkanas, P., and Kotjabopoulou, E. (2001), Quantitative sourcing of slackwater deposits at Boila Rockshelter: A record of Lateglacial flooding and Palaeolithic settlement in the Pindus Mountains, Northwest Greece. Geoarchaeology: An International Journal 16/5: 501–36. Macklin, M. G., and Smith, G. R. (2004), Pleistocene glaciation in the mountains of Greece, in J. Ehlers and P. L. Gibbard (eds.), Quaternary Glaciations—Extent and Chronology, i. Europe Elsevier, Amsterdam, 155–73. Hamlin, R. H. B., Macklin, M. G., Hughes, P. D., and Lewin, J. (2008), Glacial activity and catchment dynamics in northwest Greece: Long-term river behaviour and the slackwater sediment record for the last glacial to interglacial transition. Geomorphology 101: 44–67.

This chapter should be cited as follows Lewin, J. and Woodward, J. C. (2009), Karst geomorphology and environmental change, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 287–317.

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11

River Systems and Environmental Change Mark Macklin and Jamie Woodward

Introduction Linking river behaviour and drainage basin evolution to Quaternary environmental change, most notably the effects of climatic variability, tectonics, and human activity on runoff and sediment delivery, has a long history of research in the Mediterranean areas of Europe, North Africa, and the Near East. This field of research was initially stimulated by the (re)discovery at the beginning of the twentieth century of many Classical Period remains buried by river alluvium; perhaps the best known of which is the site of Olympia in western Greece (Huntington 1910) (Figure 11.1). The widespread evidence for large-scale shifts in river channel positions and the rapid growth of deltas and coastal alluvial plains in historical times (Judson 1963; Raphael 1973; Kraft et al. 1980; and Chapter 13) also provided much impetus for this research. In addition, archaeological investigations carried out soon after the Second World War in Algeria (Gaucher 1947), Italy (Selli 1962), Libya (McBurney and Hey 1955) and Spain (Gigout 1959) resulted in the recovery of large numbers of Palaeolithic stone tools from Pleistocene fluvial deposits. These early examples of what has now become more widely known as ‘geoarchaeology’ (Davidson and Shackley 1976; Butzer 1977) or ‘alluvial archaeology’ (Macklin and Needham 1992) were, with their strong interdisciplinary focus, highly innovative and ahead of their time in the way they integrated archaeology, geomorphology, and geochronology. Building on this theme, the principal aim of this chapter is to consider how river systems in the Mediterranean

region have responded to the environmental changes that took place during the Late Quaternary–a time interval corresponding approximately to the last 130,000 years. There are a number of reasons for choosing this period for reviewing river-environment interactions in the Mediterranean: 1. It encompasses the last glacial–interglacial cycle (c.130 to 10 ka) for which there is now abundant global evidence from polar ice cores, speleothem records, and lake and marine sediments, for both longand short-term changes in climate. These changes included massive reorganizations of the atmosphereocean-cryosphere systems—often over timescales of less than 100 years (Lowe and Walker 1997)—and they are clearly recorded in the Mediterranean region (see Allen et al. 1999 and Chapter 4). These climatic oscillations (characterized by episodes of warm and cold, wet and dry conditions) had major effects on river behaviour— both indirectly, through changes in vegetation cover (one of the primary controls of river basin hydrology and sediment supply) and directly, through changing atmospheric circulation and associated weather systems that give rise to extreme hydrological events, such as floods and droughts (Greenbaum et al. 2006). 2. In the more tectonically active parts of the Mediterranean, such as south-east Spain and much of Greece and Turkey (see Vita-Finzi 1986, and Chapter 1), the Late Quaternary Period is of sufficiently long duration to allow observation of the impact of endogenic processes (including crustal movements and volcanic activity) on river network configuration and rates of channel incision—as well as on long-term patterns of fluvial

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Fig. 11.1. A building under excavation at the archaeological site of Olympia in the valley of the Alfíos River in the western Peleponnese, Greece. This is the ancient site of the Olympic Games and the entire site was buried by alluvium from the Kladeos River which is a small right bank tributary of the Alfíos. The doorway on the right and the room on the left are still filled with alluvium (photo: Jamie Woodward).

sedimentation and storage (Harvey and Wells 1987; Collier et al. 1995; Westaway et al. 2004). Tectonic activity is also an important control on the location and development of badlands as discussed in Chapter 1. 3. This is a key period for the development of our own species. Anatomically modern humans entered Europe for the first time via the eastern Mediterranean around 100,000 years ago and, by about 30 to 25 ka, Neanderthals had disappeared from the archaeological record. The long-term dispersal and seasonal subsistence patterns of these early humans were commonly focused on coastal environments and in river valleys. Thus, many important Middle and Upper Palaeolithic sites are found in fluvial settings (e.g. Higgs and Vita-Finzi 1966; Gamble 1986; Bailey 1997) and the reconstruction of Late Quaternary river environments is often key to their interpretation (Pope and van Andel 1984; Bailey et al. 1990; Macklin et al. 1997).

4. The Late Quaternary includes the present Holocene interglacial during which human activity has had an increasing impact on the Mediterranean environment (Chapter 9) including its river systems. Indeed, one of the most challenging and long-standing problems that confronts the study of Holocene river development in the region is how to isolate the effects of human activity on the nature and rate of fluvial erosion and deposition (Vita-Finzi 1969; Butzer 1980; Lewin et al. 1995; Bintliff 2002). Over the past decade or so, this has been a major research focus facilitated by the increasing availability of continuous land- and ocean-based proxy-climate records that document significant Holocene climatic variability in the Mediterranean region extending into the historical period (e.g. Casford et al. 2001; Sadori 2001; Roberts et al. 2001). 5. In order to identify and understand both the processes that govern river behaviour and causality

Rivers and Environmental Change

relationships in fluvial systems, it is essential that alluvial units are reliably dated. Increasingly robust fluvial geochronologies based on radiometric dating methods such as luminescence and uranium series are now available for a growing number of Mediterranean river basins and these techniques have led to new insights into Middle and Late Pleistocene river behaviour (Macklin et al. 2002; Woodward et al. 2008). In addition, the development of accelerator mass spectrometry has revolutionized the dating of terminal Pleistocene and Holocene age fluvial deposits by the virtue of being able to date very small samples of organic material that could not have been dated using conventional radiocarbon dating methods (Woodward et al. 2001; Thorndycraft and Benito 2006). This has produced a rapidly expanding database of radiometrically dated Late Glacial and Holocene fluvial units that, in conjunction with new high-resolution proxy climate records, have enabled the relationship between river behaviour and climate change to be documented and interrogated in increasing detail over both space and time (e.g. Abbott and Valestro 1995; Woodward et al. 2001; Macklin et al. 2002; Thorndycraft and Benito 2006; Zielhofer et al. 2004). This chapter is divided into five sections. The first is a review of Mediterranean river systems and related environmental change research that highlights some of the major themes that have emerged over the last five decades or so. The second section presents a new typology of Late Quaternary Mediterranean river development, with particular reference to the tectonic and geological settings of river basins and the effects of Pleistocene environmental change—including glaciation. In this section, Mediterranean river responses to environmental change are examined over the last glacial– interglacial cycle (c.130 to 10 ka) where the effects of both orbitally driven and millennial-scale climatic variability are evident. The third part of this chapter examines the Holocene record (the last c.11,500 years)—where human impact and climatic oscillations were the key drivers of change within river basins. Fourthly, we examine river behaviour during the last neoglacial cycle (sensu Rumsby and Macklin 1996), which started c.1, 000 years ago and includes the socalled ‘Medieval Warm Period’ and ‘Little Ice Age’ (LIA). The final section explores direct human modifications to river channel systems that have been especially prominent over the last 150 years or so and especially in the decades since the Second World War. Chapter 8 focuses on the present-day hydrology, river regimes, and sediment load of Mediterranean catchments and

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includes a discussion of the implications of the Water Framework Directive for river systems in the European Mediterranean.

Models of Mediterranean Quaternary River Development Research conducted on Late Quaternary river histories in the Mediterranean up to the end of the 1960s was carefully reviewed and synthesized by Claudio Vita-Finzi in his acclaimed book, The Mediterranean Valleys: Geological Changes in Historical Times, published in 1969. Vita-Finzi’s study was particularly important because it was firmly grounded on almost a decade of fieldwork around the entire Mediterranean basin, extending from south-west Spain to the Jordan Valley and the Dead Sea. Using stone tools, pottery, coins, and other archaeological materials incorporated within alluvial sequences, for dating control, together with a very limited number of radiocarbon dates (five in total), he constructed a correlation scheme for Late Pleistocene (50 to 10 ka) and historical (post-Roman) alluvial units in the Mediterranean that he termed the ‘Older’ and ‘Younger’ Fills, respectively. Figure 11.2a shows a schematic crosssection of a valley in Epirus, north-west Greece where Vita-Finzi worked with the Cambridge archaeologist Eric Higgs in the 1960s (Higgs and Vita-Finzi 1966). This diagram developed from work on the Pleistocene and Holocene alluvial record in the Louros Valley near the Palaeolithic rockshelter site of Asprochaliko and it was instrumental in the development of Vita-Finzi’s Mediterranean-wide model published in 1969. It shows a prominent high terrace surface formed in a range of Pleistocene sediments—including strongly-weathered fan sediments or red beds—as well as trunk stream coarse-grained alluvial sediments that commonly interdigitated towards the valley sides with coarse limestone screes. Tufa deposits (Chapter 10) and cemented screes were common in this Older Fill. The historical or Younger Fill is set within the Pleistocene sediments in a narrow valley-floor setting created by Holocene incision and is dominated by fine-grained sediments that show little or no evidence of alteration by weathering. This phase of predominantly fluvial sedimentation has a much less variable sedimentology than the Older Fill and, according to the Vita-Finzi model, it was deposited after the Roman Period following a prolonged phase of early and mid-Holocene incision (Figure 11.2b). The thick deposits of fine-grained alluvium from the Kladeos River, that buried the ancient site of the Olympic Games at Olympia in the lower Alfios Valley of the Peloponnese,

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Mark Macklin and Jamie Woodward metres

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Fig. 11.2. (a) The Quaternary sediments of the Louros Valley in Epirus, north-west Greece (modified from Higgs and Vita-Finzi 1966). The high fan terrace is the surface of the Older Fill. These Pleistocene sediments are commonly strongly weathered and were also referred to as ‘Red Beds’. The historical fill shown in the diagram is the Younger Fill (see the main text for fuller descriptions). (b) The temporal record of channel and floodplain deposition and incision associated with the Older and Younger Fill model put forward by Vita-Finzi (1969).

are a classic example of Vita-Finzi’s Younger Fill (Figure 11.1). Vita-Finzi’s Older and Younger Fill model is probably best remembered for reducing alluvial and related colluvial deposits in Mediterranean river systems into two timed-confined eras of valley floor aggradation (see Bintliff 2002, for an insightful recent review of VitaFinzi’s contributions to Mediterranean geoarchaeology). It is clear, however, on close reading of The Mediterranean Valleys that he did in fact recognize, on the basis of the preservation of multiple alluvial terraces associated with

both the Younger (Vita-Finzi 1969: 91) and Older (ibid. 92) Fills, that each comprised several erosion and sedimentation cycles. The principal reason why Late Pleistocene and historical age alluvial units were conflated by Vita-Finzi, and by many subsequent workers, into single phases was because of the poor precision and accuracy of the dating control available at that time. Indeed this work preceded, by at least two decades, the development and application of appropriate geochronological tools (other than radiocarbon) for dating Late Quaternary alluvial

Rivers and Environmental Change

sequences. With hindsight, attributing both aggradation episodes (particularly the Younger Fill) solely to climate change seemed, even in the late 1960s, to stretch the empirical evidence, especially as high-resolution records of Late Pleistocene and Holocene climate change were not available at the time, and most reconstructions of human activity in the Mediterranean region were also not well constrained either spatially or temporally. VitaFinzi’s book was the first synthesis of Mediterranean alluvial geochronology and it set up a series of competing hypotheses of likely causative factors controlling longerterm river behaviour. This is arguably its key legacy as these bold ideas structured subsequent research and debate not only in the Mediterranean (Lewin et al. 1995; van Andel et al. 1990; Bintliff 2002) but also in the rest of Europe (Macklin et al. 2006a). There have been many advances in the four decades since Vita-Finzi’s model was first published. During the 1970s, and especially in the 1980s, a number of research programmes both in the Greek Islands (e.g. Renfrew and Wagstaff 1982) and in mainland Greece (Paepe et al. 1980; Pope and van Andel 1984; Bintliff and Snodgrass 1985), brought together interdisciplinary teams (including geomorphologists and soil scientists) coordinated by archaeologists to undertake regional Holocene landscape history projects that produced new and, in some cases, highly detailed data on human settlement in both the prehistoric and historical periods. Such information had been lacking in Vita Finzi’s and other earlier studies (e.g. Judson 1963; Raphael 1973), but these later investigations were themselves similarly limited as alluvial units were still not precisely dated and regional Holocene climate records against which supposed causative factors could be tested were also still not available (Bintliff 1992; Macklin 1995). From the mid-1980s, however, primarily as the result of the development and application of new dating techniques such as luminescence, uranium-series, and AMS radiocarbon dating, and through the reconstruction of continuous, high-resolution records of Late Pleistocene climate change (e.g. Allen et al. 1999), the chronology and controls of Pleistocene age alluvial units in the region were reassessed and in some cases radically revised (Pope and van Andel 1984; Bailey et al. 1990; Lewin et al. 1991; Fuller et al. 1998; Rose and Meng 1999; Kelly et al. 2000; Rowan et al. 2000; Woodward et al. 2001). These studies demonstrated there were multiple ‘Older Fill’ type Pleistocene alluvial units in most Mediterranean catchments and that Mediterranean rivers had been particularly responsive to short-term (millennial-side) climate change during

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the Last Glacial period. More recently, there has been a growing interest in the incidence, cause, and geomorphic impact of large flood events in the Mediterranean region and the relationship between climate change and flood frequency and magnitude. Very large boulders (>2 m) on the channel bed of the Voidomatis River in the Vikos Gorge of north-west Greece are shown in Figure 11.3. It is clear that high magnitude flood flows with very high critical shear stresses are required to move such particles and direct observations of rare catastrophic flood flows in the Mediterranean (e.g. Piegay and Bravard 1997) have fostered an increased interest in their potential role in Pleistocene and Holocene sediment fluxes and valley floor development. Studies of the occurrence and impact of large floods in Mediterranean river catchments have been undertaken on Late Pleistocene (Woodward et al. 2001; Greenbaum et al. 2006), Holocene (Benito et al. 2003a ) and historical (Barriendos Vallve and Martin-Vide 1998) timescales and particularly detailed reconstructions are beginning to emerge for the Late Holocene (Greenbaum et al. 2000; Benito et al. 2003b) and historical (Maas and Macklin 2002) periods. Over the last decade a number of our own doctoral research students have made important contributions in this area with field investigations conducted in the Peloponnese (Christopoulos 1998; Zorzou 2004), Crete, (Maas 1998; Noble 2004), north-west Greece (Hamlin 2000), and Corsica (Hewitt 2002).

Towards a Late Quaternary Evolutionary Typology of Mediterranean River Basins The present-day morphology of river basins that drain to the Mediterranean Sea, as well as the nature and configuration of Late Quaternary fluvial sedimentary sequences found within them, results from the interplay between four major series of relief-forming factors (Macklin et al. 1995; Collier et al. 1995). These are crustal mobility (directed in both horizontal and vertical directions), rock type, periodic climate, and sea-level change and, in more recent times, human action (Chapter 1). The effects of short- and long-term climate change and anthropogenic activity will be discussed later in this chapter. In this section the roles of tectonics and structural controls on Late Quaternary river development are considered, together with the influence (directly and indirectly) of Pleistocene glaciation on fluvial regime, catchment sediment supply, and river sediment fluxes. Glaciation was particularly important in the mountain catchments of the northern part of the Mediterranean (Chapter 12; Hughes et al. 2006a). Taking into

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Fig. 11.3. Large boulders (>1.5 m) on the bed of the channel of the Voidomatis River in the Vikos Gorge of north-west Greece. Mediterranean mountain river systems commonly display good slope-channel coupling and the steep scree slopes in the background deliver coarse sediment directly to the channel system (photo: Jamie Woodward).

consideration tectonic-setting, in conjunction with the presence or absence of Late Pleistocene (or earlier) glaciation in a catchment, a simple evolutionary typology of Late Quaternary fluvial systems in the Mediterranean basin is outlined.

Tectonic Setting and River Basins The Mediterranean basin forms the boundary zone between the Eurasian, African, and Arabian Plates (Chapter 1). The interaction of these plates has produced the Alpine fold belt that extends across the Mediterranean from Gibraltar to the Middle East. The region has an extremely complex and variable structure that comprises a number of microplates that have, in some cases, very different tectonic and resulting stratigraphic histories to the adjoining Eurasian and African cratons (Dewey et al. 1973). The result of this long-term collisional history, as far as Late Quaternary river basin

development and fluvial environments are concerned, has been to produce, in general terms, three rather distinct tectonic settings around the basin for river systems that drain to the Mediterranean Sea (Macklin et al. 1995): 1. The first of these is the Precambrian African plate underlying much of North Africa. With the exception of the coastal areas of Cyrenaica in northeast Libya, this is mainly a low elevation desert environment with very infrequent seismic activity (Chapter 16). In the eastern Mediterranean, however, it is diversified by rifting in the Sinai and the Jordan Valley. 2. The second is the folded and partly metamorphosed Variscan massifs of the Iberian Peninsula, Corsica, and Sardinia in the western Mediterranean. In eastern Spain, flat-lying or gently folded Mesozoic and Cenozoic sediments cover these

Rivers and Environmental Change

massifs. These areas are seismically quiescent and rarely affected by earthquakes (Chapter 16). 3. The third and largest landscape element in the Mediterranean is the Alpine fold-and-thrust belt that runs across the entire region from the Maghreb and Pyrenees in the west, through the Apennines, Sicily, and Alps proper in the central Mediterranean and extending eastward to Greece and Turkey (Chapter 1). From the point of view of river development, an important characteristic of all of these areas is their high relief, the active nature of tectonics throughout the Quaternary, and the availability in many catchments of terrains formed in mechanically weak, readily erodible lithologies such as flysch and marl (Woodward 1995). Tectonics and bedrock lithology have exerted a significant influence over fluvial systems in the Mediterranean basin because of their influence upon large-scale drainage basin morphology (size and shape) and river long-profiles and sediment fluxes—as shown for mainland Greece by Collier et al. (1995). The development of Mediterranean river systems during the Pleistocene is presented below by tectonic setting (1–3). This is followed by a Mediterranean-wide discussion of the Holocene record of river basin system response to environmental change.

Tectonic Setting 1: The Precambrian Plate of North Africa Along the presently tectonically quiescent coastline of North Africa, river systems in northern Libya that drain to the Mediterranean Sea display Late Quaternary fluvial landforms and sedimentary sequences which are generally similar over a distance approaching 1,000 km—from Tripolitania in the west (Hey 1962; Vita-Finzi 1969; Anketell et al. 1995; Gilbertson et al. 1996) to Cyrenaica in the east (Vita-Finzi 1969; Rowan et al. 2000). This suggests that, in this region, fluvial geomorphic responses to environmental change over this period have been strongly conditioned by regional geology and Pleistocene base-level change histories. Two large-scale landform elements are present in this region: a coastal plain that ranges from about 5 to 150 km in width, which is flanked to the south by a limestone plateau that rises steeply from the plain in a series of fault-guided escarpments (Figure 11.4a) lying up to 500 m above sea level. The plateau is deeply dissected by ephemeral wadi systems that emerge from the uplands onto the coastal plain as a series of coalescing alluvial

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fans. Due to progressive base-level fall since the Middle Pleistocene, these fans have a telescopically segmented form (sensu Blissenbach 1954) because the wadi channels have cut down into pre-existing fan sediments over several stages of development. The limestone uplands contain numerous inter-montane basins and these are commonly infilled with up to 30 m of Pleistocene alluvial and colluvial gravels. Schematic block diagrams of the Late Quaternary history of a typical Tripolitanian wadi produced by VitaFinzi (1969) (Figure 11.4b) show virtually identical sedimentary architectures and terrace morphologies to those recorded by Rowan et al. (2000) in Cyrenaica more than 500 km to the east. A detailed study of coastal alluvial fan development in the catchment of the Wadi Zewana, in north-eastern Libya, has been carried out in conjunction with an investigation of a series of alluvial fills within the main valley. This provided the first radiometrically dated geochronology of Middle and Late Pleistocene river activity in one of the least researched areas of the Mediterranean (Rowan et al. 2000). The Wadi Zewana drains a small (10 km2 ) catchment that meets the coast to the east of Tolmeita in Cyrenaica. The headwaters of this system lie above 350 m on the northern slopes of the Jebal Akhdar. Last interglacial bioclastic beach rock and overlying aeolianite (cemented dune sand) exposed in a coastal cliff eroding the toe of Zewana fan (Figure 11.4a) are found just above present sea level and indicate little or no uplift during the Late Pleistocene. Three major Late Pleistocene alluvial terraces have been mapped and radiometrically dated within the valley of the Wadi Zewana. The oldest and most extensive fill forms a prominent paired terrace approaching 25 m above the bed of the present river channel (Figure 11.5a). It comprises a sequence of generally flat-bedded fluvial sub-rounded gravels and sandy silts within which poorly sorted and angular colluvial gravels that thicken towards the edges of the valley are interposed (Figure 11.5b). A conspicuous feature of both fluvial and colluvial gravel units in this fill is a fine-grained matrix of terra rossa that has been eroded from the adjacent limestone slopes. This is a strongly weathered red soil found on limestone terrains across the Mediterranean (Chapter 6). Vita-Finzi (1969) noted a similarly high content of reworked terra rossa in Pleistocene wadi fills of Tripolitania. Uranium series, ESR, and OSL ages show that accelerated slope erosion, soil and regolith stripping, and valley floor infilling began sometime between c.80 and 70 ka and ended before c.42 ka when the river trenched back down to bedrock before refilling the valley floor. To establish the broader context of this record, these ages are shown

(i)

(ii)

(vii)

(iii)

(iv)

(v)

(vi)

Fig. 11.4. (a) View looking across the coastal plain in Cyrenaica (north-east Libya) showing the lower course of the Wadi Zewana in the foreground (photo: Mark Macklin). (b) Block diagrams showing the evolution of the Late Pleistocene and Holocene alluvial stratigraphy at a trunk stream tributary confluence in wadi systems in Libya (modified from Vita-Finzi 1969). The geomorphological and hydrological characteristics of these Libyan wadis are typical of many wadi systems that drain the coastal regions around the south-eastern Mediterranean, including the northern Sinai and the southern Levant.

Rivers and Environmental Change

327

(a)

(b)

Fig. 11.5. (a) The lower course of the Wadi Zewana showing a c.25 m thick exposure in Late Pleistocene alluvium. (b) A section showing the Late Pleistocene river sediments with coarse-grained angular gravels exposed at the base (photos: Mark Macklin).

on Figure 11.6 along with dated alluvial units from 16 other river basins across the Mediterranean and including key study areas discussed in this chapter. Significant slope erosion in the Zewana catchment occurred at c.23 ka and this was probably the precursor of, or perhaps coeval with, a phase of sedimentation on the coastal plain that had ended around 18 ka (Figure 11.6). The third major alluvial terrace within the Wadi Zewana rises to 12 m above the present channel bed, and a date of c.12.5 ka from the middle part of the unit, and a buried Roman cross-wadi wall c.3 m below the surface of the terrace, indicates sedimentation from the end of the Late Pleistocene until incision in the post-Roman period followed by limited valley floor refilling. Although Vita-Finzi did not have radiometric dates either for the Late Glacial or for the pre-Roman Holocene in north-west Libya, as with the earlier Late Pleistocene alluvial record, there are strong parallels between the evolution of wadi systems in Tripolitania and Cyrenaica in this period. The Wadi Zewana record suggests that Late Quaternary valley floor development in the tectonically less active parts of North Africa was characterized by an unprecedented phase of regolith and bedrock erosion that took place around 69 ka. This stripped last interglacial and earlier weathered material (terra rossa) from hillslopes to leave the bare limestone uplands of today. Although the precise duration of this phase of erosion has not yet been determined, it is clear that as the bedrock became progressively exposed in the Zewana catchment, this material was increasingly supplemented by mechanically weathered limestone gravel, resulting in aggradation along the entire trunk stream as well

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Mark Macklin and Jamie Woodward Age (ka) 10

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Guadalope, Spain, 41°00'N Guadalope, Spain, 40°50'N Bergantes, Spain, 40°45'N Bergantes, Spain, 40°40'N Voidomatis, Greece, 40°40'N Torrente d’es Coco, Mallorca, 39°40'N Loutro, Greece, 37°25'N Río Aguas, Spain, 37°10'N River Evrotas, Greece, 37°00'N Omalos, Crete, 35°19'N Samaria, Crete, 35°17'N

IRSL/OSL river terrace

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Fig. 11.6. Dated alluvial units in river systems across the Mediterranean region between c.130 and 10 ka shown in association with two proxy climate records. The river basins and reaches are ordered by latitude from north to south. The legend shows the dating methods that have been employed in each case and the error bars for each date are also shown. The dated alluvial units come from papers cited in this chapter and the proxy climate records are from Tzedakis et al. (2002).

Rivers and Environmental Change

as in the larger tributaries (Figure 11.5). Following this major phase of aggradation, from c.42 ka the Wadi Zewana became progressively supply-limited with later phases of valley floor sedimentation resulting from the erosion and ‘cannibalization’ of the c.69 ka alluvial fill (see Lewin and Macklin 2003). The period shortly after c.69 ka along the southern Mediterranean littoral in Libya was a time of marked geomorphic change. Indeed, it represents probably the most significant landscape transformation that has occurred in this region for the last 100,000 years and perhaps longer (Figure 11.6). This period lies close to the Marine Isotope Stage (MIS) 5/4 transition which is dated to c.70 ka on the SPECMAP timescale and to about 75 to 80 ka in the GRIP ice core (Dansgaard et al. 1993 and Figure 11.6). It was marked by a sharp decline in sea-surface temperatures in the North Atlantic by as much as 5–6 o C (Bond et al. 1993) and the replacement of mixed woodland in southern Europe by open steppe (Tzedakis et al. 2002 and Chapter 4). This geomorphic ‘event’ set the boundary conditions for all subsequent phases of valley floor erosion and sedimentation in the region, including those that occurred during the Holocene and historical periods. It is clear that a major threshold was crossed at around 70 ka in terms of both the rates and types of hydrological and geomorphological processes operating within the river basins of the region. Before 70 ka geomorphic activity was relatively subdued in the catchment. After this date, however, there was a step-change increase in slope erosion and concomitant valley floor aggradation that was probably triggered by changes in vegetation type and cover—themselves related to altered precipitation regimes. Thereafter the Zewana basin became more responsive to a series of major climatic fluctuations at c.42, 23–18, and 12.5 ka that are recorded in marine, ice core, and pollen records (Figure 11.7).

Tectonic Setting II: The Hercynian Massifs and Foldbelts The Hercynian massifs and fold belts that form much of the central and eastern part of the Iberian Peninsula, together with the Hercynian massifs of Corsica and Sardinia that were involved in Alpine folding, comprise the second tectonic terrain with respect to a morpho-structural classification of Mediterranean river basins and their Late Quaternary development. Located away from the main plate boundaries, these areas have had relatively little seismic activity over the last 130,000 years (Chapter 16). However, gradual long-term regional uplift has created high basin

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relief (34% of peninsular Spain lies between 800 and 2,000 m above sea level) and, with extensive outcrops of sedimentary rock susceptible to mechanical breakdown and erosion, these landscapes provide ideal conditions for the formation and preservation of river terraces (Macklin and Passmore 1995; Santisteban and Schulte 2007). Glacial moraines and Pleistocene river terraces are well preserved in the mountains of Corsica (Conchon 1986), but these records are not yet securely dated. One of the most studied and securely dated Quaternary river terrace sequences in the Mediterranean is that of the River Guadalope and its major tributary the River Bergantes in north-east Spain that drain the Iberian Cordillera and flow northwards into the Ebro basin (Macklin and Passmore 1995; Fuller et al. 1996, 1998; López-Avilés et al. 1998). High resolution geomorphological mapping in the lower part of the Bergantes Valley has identified five major river terraces at 20–25 m, 16 m, 10–12 m, 5–8 m, and 2–3 m above present river level. Luminescence dating of alluvial fills demonstrates trunk river aggradation at c.110 ka, c.35–40 ka, c.25– 7 ka, c.10–13 ka, c.7–8 ka, and c.3 ka, respectively. In broad terms, the scale of Late Pleistocene fluvial aggradation events varies considerably—with the oldest (c.110 ka) being the largest and with each subsequent event becoming progressively smaller. It is evident that differences in both the scale and pattern of Late Pleistocene fluvial sedimentation were controlled primarily by the degree of coupling between tributaries and the axial river and the amount of sediment supplied from local slopes and tributary catchments. The period around 110 ka saw up to 20 m of trunk river aggradation in the Bergantes Valley associated with exceptionally high rates of erosion and sediment delivery during the cold episode of Marine Isotope Substage 5d (Figure 11.6). This represents a ‘tipping point’ in catchment evolution and Late Quaternary river development in the region and appears to be directly analogous to the exceptional phase of drainage basin transformation at c.70 ka recorded in North African Mediterranean river systems discussed above. Both the c.110 ka and 70 ka phases of accelerated geomorphic activity coincided with forest decline and the development of steppe-dominated landscapes as shown by pollen records in southern Europe (e.g. Allen et al. 1999; Tzedakis et al. 2002) (Figure 11.6). Similar vegetation changes are evident in later stadial events, but the magnitude of the geomorphic response in north-east Spain during Marine Isotope Substage 5d suggests that climatic forcing of catchment erosion, sediment yields, and fluvial system dynamics were amplified in comparison to later cold

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Dating control and depositional environment

Omalos, Crete, 35°19'N

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Alluviation Incision

Rapanas, Crete, 35°15'N Aradena, Crete, 35°13'N Oued es Seffaia, Tunisia, 34°00'N Wadi Zewana, Libya, 32°30'N Tissint feija, Morocco, 29°55'N Seyad feija, Morocco, 29°04'N - 34 - 36

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Fig. 11.7. Dated alluvial units in river systems across the Mediterranean region between c.65 and 10 ka. The shaded columns show Heinrich Events 1–6 (see Chapter 4). The river basins and reaches are ordered by latitude from north to south. The legend shows the dating methods that have been employed in each case and the error bars for each date are also shown. The dated alluvial units come from papers cited in this chapter and the proxy climate records are from Roucoux et al. (2005). The five palynological sections (a to e) from Roucoux et al. (2005) are also shown.

Rivers and Environmental Change

periods, perhaps through a combination of exceptional rates of runoff and sediment supply. Subsequent phases of valley floor refilling in northeast Spain are coeval with Heinrich event 4 (c.36–41 ka) and the Younger Dryas Stadial (Figure 11.7). Heinrich event 4 is believed to have been one of the most extreme climatic events of MIS 3, judged by the major flux of icerafted detritus in the North Atlantic from the Laurentide ice sheet (McManus et al. 1998) and the reduction in the thermohaline circulation (Tzedakis et al. 2002). In comparison to substage 5d, however, there was limited sediment supply from tributary streams during these periods and material that was delivered from tributary catchments was generally fine-grained in calibre. As elsewhere in the Mediterranean, while phases of large-scale Late Pleistocene river aggradation in the Bergantes Valley are correlated with Heinrich events and with stades of the Dansgaard-Oeschger cycles (Macklin et al. 2002), the record of river response to sub-orbital-scale climate change is more complex and spatially variable. In the tectonically quiescent Mediterranean regions of North Africa and the Iberian Peninsula, environmental changes were characterized primarily by reorganizations within the fluvial sediment system—the most notable of which were major fluctuations in tributary stream sediment delivery, and changes in alluvial storage within high-order valleys. These adjustments occurred in both areas over extended periods. In the drainage basins in north-east Spain, these began during the transition to the last cold stage (MIS 5d to 5a) and in north-east Libya were initiated during the MIS 5/4 transition (Figure 11.6). Over timescales of 104 years and longer, sediment yields have therefore been highly variable and data from the River Bergantes demonstrate very clearly that some of the larger (>1200 km2 ) Mediterranean river systems have been highly sensitive to changes in climate.

Tectonic Setting III: The Alpine Mountain Belt The largest and most geologically complex tectonic terrain in the Mediterranean is that resulting from Alpine orogenesis. It comprises the mountain chains of the Pyrenees, the Alps, the Balkans, and Turkey related to continental collision, and fold-and-thrust belts in south-east Spain, the Maghreb, and the Italian Peninsula related to back arc opening (Chapter 1). These landscapes include the most seismically active areas of the Mediterranean (Chapter 16). In many mountain regions within this tectonic province, particularly along the northern Mediterranean littoral of Italy, the

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former Yugoslavia, Greece, and Turkey, climatic cooling during the Pleistocene provided favourable conditions for glacier development (Chapter 12). Glaciers had a significant impact on sediment supply and catchment hydrology, and these effects have been observed in the Pleistocene fluvial record in river valleys draining glaciated drainage basins in north-west Greece (Lewin et al. 1991; Woodward et al. 1992, 1995 2008), northwest Slovenia (Bavec et al. 2004) and north-east Spain (Sancho et al. 2003). It is therefore logical for the purposes of developing an evolutionary typology to subdivide Mediterranean river basins within the area affected by Alpine orogenesis into those catchments which supported important Pleistocene cirque and valley glaciers and those that did not. The Pleistocene glacial history of the Mediterranean mountains has recently been reviewed by Hughes et al. (2006a ) and is examined in detail in Chapter 12.

Glaciated River Basins in the Alpine Fold Belt The widespread occurrence of glacial landforms and sediments in the headwaters of many Mediterranean river basins has been known for well over a century (Cviji´c 1900; Niculescu 1915; Almagià 1918; Chapter 12). However, outside north-east Spain (Pyrenees; Sancho et al. 2003), north-west Slovenia (South Julian Alps; Bavec et al. 2004) and, most notably, the Voidomatis River basin in the Pindus Mountains of northwest Greece (Bailey et al. 1990; Lewin et al. 1991; Woodward et al. 1992, 1995 2001; Macklin et al. 1997, 2002; Hamlin 2000; Hamlin et al. 2000; Woodward et al. 2008), the downstream impact of this glacial activity upon Pleistocene river development is still relatively poorly known. A substantial body of geomorphological, sedimentological, and geochronological data has been assembled for the Middle and Late Pleistocene history of the Voidomatis River basin (384 km2 ). A large area of the basin’s headwaters lies in the high karst of the Pindus Mountains (Mount Tymphi, 2,470 m) and was subject to the build-up and decay of cirque and valley glaciers on at least three occasions over the last 500,000 years or so (Woodward et al. 2004; Hughes et al. 2006b). The glacial sequence on Mount Tymphi is now one of the best-dated in Europe and is discussed more fully in Chapter 12. In brief, the record comprises extensive Middle Pleistocene glacial sediments and landforms that have been dated to MIS 12 (between c.478 and 423 ka) and MIS 6 (between c.190 and 126 ka) using uraniumseries methods. In contrast, only small cirque and valley glaciers and rock glaciers developed during the last

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Pl e i s t o c e n e

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TL 28.2 ± 7 ka

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U/Th 74 ± 6 ka

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ESR 24.3 ± 2.6 ka 25.0 ± 0.5 ka 26.0 ± 1.9 ka

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Fig. 11.8. (a) The Pleistocene and Holocene fluvial stratigraphy in the middle and lower reaches of the Voidomatis River basin showing eight alluvial units including the modern channel complex. Based on a figure in Hamlin et al. (2000) using data from Lewin et al. 1991; Macklin et al. 1997; Hamlin et al. 2000). Key to dating methods: U/Th = uranium thorium, ESR = electron spin resonance, TL = thermoluminescence (both dates on the U6 and U3 soils are TL dates. U2 has been dated using radiocarbon). The photographs show (b) Unit 6 in the Lower Vikos Gorge at the location of soil profile A of Woodward et al. (1994) and (c) Unit 2 in the southern part of the Konitsa basin a few hundred metres downstream of Boila rockshelter (photos: Jamie Woodward).

cold stage or MIS 5d to 2 (between c.111 and 11.5 ka) (Hughes et al. 2006b). Late Pleistocene river sediments are well preserved in the lower reaches of the Voidomatis River and a detailed record of river aggradation and dissection, and its relation to headwater glaciation, has been reconstructed (Lewin et al. 1991; Macklin et al. 1997; Woodward et al. 2001, 2008). The fluvial record in the lower reaches of the Voidomatis River is shown schematically in Figure 11.8a along with the array

of dates that constrain the phases of fluvial sedimentation. Uniquely, the geochronology combines these four methods and sample types: uranium-series dating of calcite cements, thermoluminescence dating of fine-grained alluvial soils, electron spin resonance dating of fossil teeth, and AMS radiocarbon dating of charcoal. Alluvial units (U7 to U4) have a very similar lithological composition (Figure 11.8b)—they are dominated by coarse limestone gravels (>94%) and they contain a fine-grained matrix derived from the

Rivers and Environmental Change

glacial deposits in the catchment headwaters (Woodward et al. 1992; Hamlin et al. 2000). These four units represent the Aristi Unit sediments originally described by Lewin et al. (1991) that have been related to glacial activity in the catchment headwaters (Figure 11.8b). The Aristi Unit was later subdivided into four phases of aggradation and incision following detailed mapping of the lower reaches of the basin in the late 1990s and, crucially, through the application of uraniumseries dating to carbonate cements in the coarse gravel matrix (Hamlin et al. 2000) (Figure 11.8a). These ages are also plotted on Figures 11.6 and 11.7 along with other radiometrically dated alluvial units from across the Mediterranean. These phases of aggradation represent periods of enhanced sediment supply to the lower reaches from a range of sediment sources including: direct input from headwater glaciers, fluvial reworking of pre-existing glacial and alluvial sediments, and cold climate weathering of hillslopes. It is interesting to note that the Voidomatis River terrace sequence provides a more detailed archive of Late Pleistocene environmental change than does the glacial record in the catchment headwaters (Chapter 12). Unit U8 is the Kipi Unit of Lewin et al. (1991) and it has a very different lithological composition to the later, limestone-rich units. It is dominated by flysch and ophiolite gravels (Figure 11.8b) and has a flysch-rich matrix. It must therefore predate the first phase of glaciation in the catchment (MIS 12) and the TL date of >150 ka (Figure 11.8a) is a minimum age. U3 is the youngest Pleistocene alluvial unit and is equivalent to the Vikos Unit of Lewin et al. (1991). The lithology of this unit shows that it was deposited after the final period of Late Pleistocene glacial activity on Mount Tymphi and it overlaps with a phase of Late Glacial slackwater sedimentation recorded in the deposits of Boila rockshelter in the lower reaches of the Voidomatis River (Woodward et al. 2001) and shown in Figure 11.9. The postglacial record of river behaviour is dominated by progressive incision that is interrupted in the late Holocene by the development of a distinctive low terrace (U2) formed in fine-grained overbank sediments (Figure 11.8c).

Non-glaciated River Basins in the Alpine Fold Belt The Sorbas basin, part of the Betic Cordillera of southeast Spain, lies within a rapidly uplifting area, with Pliocene to recent average uplift rates of c.80 m Ma−1 (Mather 2000; Chapter 1). This basin was not glaciated during the Pleistocene and provides one of the bestdocumented examples of the effects of tectonics on Late Pleistocene and earlier drainage network development in the Mediterranean (Harvey and Wells 1987).

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Fig. 11.9. The sediments exposed during the excavations at Boila rockshelter in the lower reaches of the Voidomatis River in northwest Greece (after Woodward et al. 2001). The bottom of the section shows the fine-grained slackwater sediments produced by large floods prior to the Late Upper Palaeolithic use of the site. The sediments above contain angular limestone clasts derived from the walls and ceiling of the rockshelter. The scale is in cm (photo: Jamie Woodward).

Accelerated headward erosion of the Río Aguas, stimulated by regional differential uplift, resulted in progressive reorganization of the drainage network by river capture. The most dramatic modification of the network was a major capture of the south-flowing proto Aguas/Feos by the east-flowing lower Aguas, an aggressive subsequent stream that was developing by headward erosion along the outcrop of a weak marl unit. Three pre-capture river terraces can be traced following the proto-drainage across the Sorbas basin and through the southern mountain ranges (Harvey and Wells 1987). Uranium series dating of pedogenic calcretes developed on the surface of these river terraces

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Fig. 11.10. The deeply incised valley floor and Quaternary terraces of the Río Aguas in south-east Spain (see Chapter 1). This reach is a few kilometres downstream of Sorbas and shows the Urra field centre on the high terrace (photo: Jamie Woodward).

has recently been undertaken (Candy et al. 2005). These workers argue that the terraces date to c.304 ka, 207 ka, and 70 ka, respectively, with capture of the Río Feos by the lower Aguas taking place some time between 69.8 and 67.9 ka. These dates, along with three further OSL dates published by Schulte et al. (2007) are also plotted on Figure 11.6. The dates for the two oldest terraces indicate that river incision of 10–15 m began towards the end of the interglacials of MIS 7 and 9, followed by limited refilling (c.5 m) of the valley floor. In the case of the c.70 ka river terrace, calcite-cemented root mats in the middle part of the gravel-dominated unit show that it comprises two aggradational phases; one dated to sometime before 77 ka and the second shortly after 77 ka (Candy et al. 2005) (Figure 11.6). Post-capture terraces (5.25 m at Arles and this equates to a discharge of >7,000 m3 s−1 . These big floods were more frequent in the decades from 1651 to 1720 and 1751 to 1860. This flood-dominated regime was interrupted by periods of low flood frequency from 1721 to 1750 and from 1861 to 1995. The frequency of these large floods has fallen since the end of the LIA with eight to nine per decade between 1850 and 1900, four to five per decade between 1900 and 1950 and only two to three per decade between 1950 and 2000 (Pichard 1995; Arnoud-Fassetta 2003). Further south in the western Mediterranean, Hewitt (2002) has developed a flood history for the Figarella River (132 km2 ) that drains to Calvi Bay from the mountains of north-west Corsica. Here, flood deposits—including boulder splays, boulder bars, and boulder berms—are well preserved across four terrace surfaces and these have also been dated using lichenometry. Hewitt (2002) identified 18 flood units

that represent ten periods of enhanced flooding over the last four centuries, with the oldest units dated to between 1572 and 1595. Most of the flood units preserved in the Figarella record were deposited by floods that took place in the nineteenth century. In common with the record of large floods from the Rhône (Pichard 1995), the Figarella River has a much lower frequency of large floods in the twentieth in comparison to the nineteenth century. Hewitt (2002) has shown that the Figarella flood history shows a good agreement with the regional-scale record of large floods compiled from a range of data sources by Barriendos Vallve and Martin-Vide (1998) for ten river catchments that drain to the coast of north-east Spain (Figure 11.18). This figure compares flood records from documentary sources with geomorphological records from steep gradient river systems located in Corisca, Crete, and north-west Greece. Broadly speaking, five periods of enhanced flooding can be identified within these records: 1. The late sixteenth century (north-east Spain, southern France, and Corsica). 2. The late seventeenth to early eighteenth century (southern France). 3. Late eighteenth century (across the Mediterranean, but not in Crete). 4. The mid and late nineteenth century (across the Mediterranean). 5. The first half of the twentieth century (Corsica, north-west Greece, and Crete). The record from north-west Greece is the only one in this data set to show enhanced flooding in the second half of the twentieth century. It is important to appreciate that the records from France and north-east Spain shown in Figure 11.18 come from documentary and instrumental sources, whilst those from Corsica, north-west Greece, and Crete are from geomorphological evidence dated by lichenometry and these three may partly reflect the influence of preservation factors. The Corsican record goes back as far as the longest historical records so far published for the western Mediterranean.

Recent Human Modifications to River Channel Systems Over the last one hundred and fifty years or so, and especially since the Second World War, river channel systems in the Mediterranean region have been subjected to a range of modifications resulting from engineering works and resource exploitation. Impacts include

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Fig. 11.18. Flood histories from five parts of the Mediterranean since AD 1500 based on geomorphological (solid shading) and documentary (hatched shading) records.

dam construction and reservoir development for hydroelectric power generation and flood control, as well as water resource development for irrigation and public water supply (Palanques et al. 1990; Surian 1999; Arnaud-Fassetta 2003). Reservoir sedimentation is now a major problem across the region (Lahlou 1988; Woodward 1995; Woodward and Foster 1997; Chapter 8) and this has led to river channel incision in downstream reaches (Surian and Rinaldi 2003). In the middle and lower reaches of many river catchments, engineers have drained wetlands, built flood embankments, straightened channels, and developed extensive water abstraction and irrigation schemes. The second half of the twentieth century has seen an increasing demand for sands and gravels for a range of civil engineering projects including road building and urban development. This has encouraged intensive gravel extraction from floodplains and from the active channel zone in many Mediterranean rivers (Nicholas et al. 1999). The response of several Italian rivers to river engineering projects and gravel extraction has been reviewed by Surian and Rinaldi (2003). They have shown that 3–4 m of channel incision is a common response in reaches of the Po, Arno, and Piave rivers and channel narrowing of more than 50 per cent has been observed. These channel responses are more rapid immediately following the disturbances and the larger rivers tend to have longer recovery times. Surian and Rinaldi (2003) propose a model for the main styles of

river channel response to recent human intervention in Italy with braided rivers adjusting through narrowing of the active channel zone with varying rates of incision, whereas single-thread river channels adjust via more pronounced vertical incision with variable degrees of channel narrowing. Precious and base metal mining in the Mediterranean has a very long history with mining activity in the Río Tinto of south-west Spain being amongst the oldest in Europe dating back more than 5,000 years. Historical mining activity has resulted in significant and widespread contamination in many Mediterranean river basins with floodplain sediment in mining-affected catchments presently acting as a major secondary source of pollution (Macklin et al. 2006b). Environmental problems also arise from present-day mining activities with the April 1998 Aznalcóllar tailings dam failure being one of the most serious river pollution incidents in Europe since the Second World War (HudsonEdwards et al. 2003; Macklin et al. 2006b; Turner et al. 2002, 2008; Turner 2004). On 25 April 1998, a tailings dam failed at the Frailes lead and zinc mine at Aznalcóllar near Seville in south-west Spain. The spill released over 4 million m3 of toxic slurries into the Río Agrio, a tributary of the Río Guadiamar. This accident covered several thousand hectares of farmland with contaminated water and sediment and has had a long-lasting impact on the Río Guadiamar that received most of the tailings waste and

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Fig. 11.19. The deeply incised channel zone of the Alfíos River in western Greece. Intensive in-stream gravel extraction has produced rapid channel degradation and bank retreat (photo: Jamie Woodward).

the internationally important wetlands of the Doñana National Park on the Atlantic coast (see Chapter 9). This UN World Heritage Area is still receiving significantly elevated concentrations of heavy metals ten years after the spill. Sediment-borne metals derived from the erosion and remobilization of contaminated river channel and floodplain sediment is the major problem at this site. During the course of the twentieth century, the sediment load reaching the Rhône delta declined by a factor of 4.7 from 35 to 7.39 million tonnes per year largely due to the construction of large dams in the catchment. Channel incision on the Rhône delta began in 1895, eventually leading to a lowering of the floodplain water table and salinization of floodplain soils as saline waters migrated upstream from the coastal zone (Arnaud-Fassetta 2003). In western Greece, the lower reaches of the Alfios River basin (3,600 km2 ) were transformed in less than three decades following intensive gravel extraction from the bed of the active channel (Figure 11.19). Between 1967 and 1995 more than 17 million m3 of gravel were extracted using large cranes and drag buckets (Christopoulos 1998; Nicholas et al. 1999). This activity took place in the reaches immediately downstream of

the Flokas Dam which is only about 5 km downstream of the archaeological site of Olympia—where this chapter began! The dam prevents bed load replenishment in the reaches downstream and this amplified the response of the channel to a major perturbation to the coarse sediment budget. This led to a period of rapid and sustained channel degradation resulting in 6–8 m of vertical incision and mean annual bank erosion rates of about 10 m per year with major bank failures associated with large flood events. Air photographs show that, in the gravel extraction reaches (immediately downstream of the Flokas Dam), the main channel of the Alfios River may have widened by up to 100 m between 1960 and 1990 (Christopoulos 1998; Nicholas et al. 1999). Alluvial gravel extraction is common throughout the Mediterranean region where gravel bed rivers with broad channel zones and seasonally low stages provide ready access to this important resource. The regulation of these gravel extraction activities is commonly ineffective and excessive rates of extraction can promote rapid channel changes that may propagate both upstream and downstream. This has led to falling water tables, elevated turbidity, and large tracts of valley floor with severely degraded channel and riparian habitats.

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Fig. 11.20. Summary of human impacts on the river channel systems in the lower reaches and delta complex of the Axíos, Aliakmon, and Gallikos rivers in north-east Greece over the last century (modified from Kapsimalis et al. 2005). The darker the bars the greater the intensity of the human impact.

Fish species that spawn in gravel substrates are also threatened by gravel extraction. Cote et al. (1999) have shown how populations of river blennies (Salaria fluviatilis), that have a wide circum-Mediterranean distribution, are threatened by gravel extraction because structural alterations to gravel beds render the habitats unsuitable for breeding. In north-east Greece, Kapsimalis et al. (2005) have investigated recent changes in the lower reaches of three large rivers systems that drain to the Inner Thermaikos Gulf in the north-west Aegean Sea. The catchments of the Axios (23,747 km2 ), Aliakmon (9,250 km2 ), and Gallikos (1,230 km2 ) rivers accounted for a mean annual water discharge of approximately 276 m3 s−1 between 1926 and 1970, but this has declined in recent years to c.130 m3 s−1 because of abstractions for irrigation and urban consumption. Figure 11.20 provides a summary of the major changes between 1850 and 2000 to the lower reaches of these rivers following the increasing intensity of human impacts on water and sediment fluxes. The list of human activities shown in Figure 11.20 is typical of the impacts seen in many river basins across the Mediterranean region in the twentieth century (see Surian and Rinaldi 2003). These impacts have led to dramatic transformations to alluvial channels, floodplains, and deltaic environments. The net result of these activities has been a radical transformation of catchment water and sediment budgets and of

channel and floodplain environments across the region and these impacts are especially apparent in the more populous European Mediterranean.

Conclusions The Mediterranean basin provides a unique region in which to examine the long-term impact of environmental change on river system dynamics. The tectonic disposition of the region means that Quaternary fluvial sediments and landforms are commonly well preserved in river basins across the region and they represent a very significant archive of environmental change. The rich archaeological records have allowed human– river environment interactions to be studied in unusual detail from Palaeolithic to recent times. The interaction between fluvial geomorphologists and archaeologists over many decades has generated important ideas and debates about river system development that have resonated well beyond the Mediterranean world. Recent advances include the development of much more robust dating frameworks for fluvial sequences and a greater appreciation of the significance of large floods in the creation of the fluvial archive. However, there is still much work to do if we are to resolve outstanding issues associated with the investigation and interpretation of the Holocene record. More detailed records

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of land use change are needed in many areas. Closer collaboration at the basin scale between geoscientists and archaeologists and historians would improve our understanding of the impacts of land use change on the behaviour of river basin systems. The flood hydrology and sediment loads of many catchments across the Mediterranean were modified significantly by the climate changes associated with the Little Ice Age and data from this period provide a valuable guide for our interpretation of earlier parts of the fluvial archive. Researchers now have a much better appreciation of the role of large flood events in Mediterranean river systems through the systematic investigation of slackwater sediments, boulder flood units, documentary evidence, and the monitored hydrological record. Environmental changes associated with direct human impacts on river systems have intensified over the last 150 years or so and many valley floor environments in the Mediterranean have been degraded by resource exploitation, urban expansion, and the impacts of water and sediment pollution.

Acknowledgements This chapter was completed in the Acropole Hotel in Khartoum about 3,000 km upstream of the Mediterranean Sea. We would like to thank the external reviewer for a thoughtful review of the text and Nick Scarle for producing the diagrams. Our work in the Mediterranean has been supported by various bodies over the last two decades including the NERC and by many interdisciplinary archaeological projects. It has also involved collaboration with several generations of Ph.D. students and with many colleagues and good friends. We are grateful to them all.

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and Needham, S. (1992), Studies in alluvial archaeology: potential and prospect, in S. Needham and M. G. Macklin (eds.), Alluvial Archaeology in Britain. Oxbow, Oxford, 9–23. and Passmore, D. G. (1995), Pleistocene environmental change in the Guadalope basin, northeast Spain: fluvial and archaeological records, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Rotterdam, Balkema, 103–113. Lewin, J., and Woodward, J. C. (1995), Quaternary fluvial systems in the Mediterranean basin, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Balkema, Rotterdam, 1–25. (1997), Quaternary River Sedimentary Sequences of the Voidomatis Basin, in G. N. Bailey (ed.), Klithi: Palaeolithic Settlement and Quaternary Landscapes in Northwest Greece, ii. Klithi in its Local and Regional Setting. McDonald Institute for Archaeological Research, Cambridge, 347–59. Fuller, I. C., Lewin, J., Maas, G. S., Passmore, D. G., Rose, J., Woodward, J. C., Black, S., Hamlin, R. H. B., and Rowan, J. S. (2002), Correlation of Late and Middle Pleistocene fluvial sequences in the Mediterranean and their relationship to climate change. Quaternary Science Reviews 21: 1633–44. Johnstone, E., and Lewin, J. (2005), Pervasive and longterm forcing of Holocene river instability and flooding by centennial-scale climate change. The Holocene 15: 937–43. Benito, G., Gregory, K. J., Johnstone, E., Lewin, J., ´ Michczynska, D. J., Soja, R., Starkel, L., and Thorndycraft, V. R. (2006a ), Past hydrological events reflected in the Holocene fluvial record of Europe. Catena 66: 145–54. Brewer, P. A., Hudson-Edwards, K. A., Bird, G., Coulthard, T. J., Dennis, I. A., Lechler, P. J., Miller, J. R., and Turner, J. N. (2006b), A geomorphological approach to the management of rivers contaminated by metal mining. Geomorphology 79: 423–47. McManus, J. F., Anderson, R. F., Broecker, W. S., Fleisher, M. Q., and Higgins, S. M. (1998), Radiometrically determined sedimentary fluxes in the sub-polar North Atlantic during the last 140,000 years. Earth and Planetary Science Letters 155: 29–43. Magny, M., Bégeot, C., Guiot, J., and Peyron, O. (2003), Contrasting patterns of hydrological changes in Europe in response to Holocene climate cooling phases. Quaternary Science Reviews 22: 1589–96. Mather, A. E. (2000), Adjustment of a drainage network to capture induced base-level change. Geomorphology 34: 271–89. Harvey, A. M., and Brencheley, P. J. (1991), Halokinetic deformation of Quaternary river terraces in the Sorbas basin, SE Spain. Zeitfrich für Geomorphologie Suppl. 82: 87–97. Moody, J. and Grove, A. T. (1990), Terraces and enclosure walls in the Cretan landscape, in S. Bottema, G. Entjes-Nieborg, and W. van Zeist (eds.), Man’s Role in the Shaping of the Eastern Mediterranean Landscape. A. A. Balkema, Rotterdam, 183–91. Nicholas, A. P., Woodward, J. C., Christopoulos, G., and Macklin, M. G. (1999), Modelling and monitoring the impact of dam construction and gravel extraction on rates of bank erosion in the Alfios River, Peloponnese, western Greece, in A. G. Brown and T. A. Quine (eds.), Fluvial Processes and Environmental Change. John Wiley & Sons, Chichester, 117–37. Niculescu, C. (1915), Sur les traces de glaciation dans le massif Smolica chaîne du Pinde méridional. Bulletin de la Section Scientifique de l’Academie Roumaine 3: 146–51.

Rivers and Environmental Change Noble, J. (2004), Response of a steepland river system to late Quaternary environmental change: south central Crete. Ph.D. Thesis, University of Wales, Aberystwyth. Paepe, R., Hatziotis, M. E., and Thorez, J. (1980), Geomorphological Evolution in the Eastern. Mediterranean Belt and Mesopotamian Plain. Report for the International Geological Correlation Programme Project 146: River flood and lake level changes. Palanques, A., Plana, F., and Maldonado, A. (1990), Recent influence of man on the Ebro margin sedimentation system, northwestern Mediterranean Sea. Marine Geology 95: 247–63. Pichard, G. (1995), Les crues sur le bas Rhône de 1500 à nos jours. Pour une histoire hydro-climatique. Méditerranée 3/4: 105–16. Piegay, H. and Bravard, J.-P. (1997), Response of a Mediterranean riparian forest to a 1 in 400 year flood, Ouveze River, Drome-Vaucluse, France. Earth Surface Processes and Landforms 22: 31–43. Pope, K. O. and van Andel, T. H. (1984), Late Quaternary alluviation and soil formation in the southern Argolid: its history, causes and archaeological implications. Journal of Archaeological Science 11: 281–306. Raphael, C. N. (1973), Late Quaternary Changes in Coastal Elis, Greece. Geographical Review 63: 73–89. Renfrew, A. C. and Wagstaff, M. (1982), An Island Polity: The Archaeology of Exploitation in Melos. Cambridge University Press, Cambridge. Roberts, N., Reed, J., Leng, M. J., Kuzucuo˘glu, C., Fontugne, M., Bertaux, J., Woldring, H., Bottema, S., Black, S., Hunt, E., and Karabıyıko˘glu, M. (2001), The tempo of Holocene climatic change in the eastern Mediterranean region: new highresolution crater-lake sediment data from central Turkey. The Holocene 11: 721–36. Rose, J. and Meng, X. (1999), River activity in small catchments over the last 140 ka, northeast Mallorca, Spain, in: A. G. Brown and T. A. Quine (eds.), Fluvial Processes and Environmental Change. Wiley, Chichester, 91–102. Roucoux, K. H., de Abreu, L., Shackleton, N. J., Tzedakis, P. C., (2005), The response of NW Iberian vegetation to North Atlantic climate oscillations during the last 65 kyr. Quaternary Science Reviews 25: 1637–53. Rowan, J. S., Black, S., Macklin, M. G., Tabner, B. J., and Dore, J. (2000), Quaternary environmental change in Cyrenaica evidenced by U-Th, ESR and OSL of coastal alluvial fan sequences. Libyan Studies 31: 5–16. Rumsby, B. A. and Macklin, M. G. (1996), River response to the last neoglacial (the ‘Little Ice Age’) in northern, western and central Europe, in K. J. Gregory (ed.), Global Continental Changes: The Context of Palaeohydrology. Geological Society Special Publication, London 115: 217–33. Sadori, L. (2001), The Postglacial record of environmental history from Lago di Pergusa, Sicily. The Holocene 11: 655–71. Sancho, C., Peña Monné, J. L., Lewis, C., McDonald, E., and Rhodes, E. (2003), Preliminary dating of glacial and fluvial deposits in the Cinca River Valley (NE Spain): chronological evidences for the Glacial Maximum in the Pyrenees? in B. Ruíz-Zapata, M. Dorado-Valiño, A. Valdemillos, M. J. Gil-García, T. Badají, I. Bustamante, I. Mendizábal (eds.), Quaternary Climatic Changes and Environmental Crises in the Mediterranean Region. Universidad de Alcalá de Henares, Madrid, 453–6. Santisteban, J. I. and Schulte, L. (2007), Fluvial networks of the Iberian Peninsula: a chronological framework. Quaternary Science Reviews 26: 2738–57.

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Schulte, L., Julià, R., Burjachs, F., and Hilgers, A. (2007), Middle Pleistocene to Holocene geochronology of the River Aguas terrace sequence (Iberian Peninsula): fluvial response to Mediterranean environmental change. Geomorphology 98: 13–33. Selli, R. (1962), Le Quaternaire marin du versant Adriatique Ionien de la péninsule italienne. Quaternaria 6: 391–413. Surian, N. (1999), Channel changes due to river regulation: the case of the Piave River, Italy. Earth Surface Processes and Landforms 24: 1135–51. and Rinaldi, M. (2003), Morphological response to river engineering and management in alluvial channels in Italy. Geomorphology 50: 307–26. Thorndycraft, V. R. and Benito, G. (2006), Late Holocene fluvial chronology of Spain: The role of climatic variability and human impact. Catena 66: 34–41. Turner, J. N. (2004), Ageomorphological-geochemical assessment of the impacts of the April 1998 Aznalcóllar tailings dam failure on the Rio Guadiamar, southwest Spain. Ph.D. Thesis, University of Wales, Aberystwyth. Brewer, P., Macklin, M. G., Hudson-Edwards, K. A., Coulthard, T. J., Howard, A. J., and Jamieson, H. E. (2002), Heavy metal and as transport under low and high flows in the River Guadiamar three years after the Aznalcollar tailings dam failure: Implications for river recovery and management, in J. M. Garcia, J. A. A. Jones, and J. Arnaez (eds.), Environmental Change and Water Sustainability. IPE-CSIC, Zaragoza, 235–51. (2008) Fluvial-controlled metal and as mobilisation, dispersal and storage in the Río Guadiamar, SW Spain and its implications for long-term contaminant fluxes to the Doñana wetlands. Science of The Total Environment 394: 144– 61. Tzedakis, P. C., Lawson, I. T., Frogley, M. R., Hewitt, G. M., and Preece, R. C. (2002), Buffered tree population changes in a Quaternary refugium: evolutionary implications. Science 297: 2044–7. van Andel, T. H. and Runnels, C. (1987), Beyond the Acropolis: A Rural Greek Past. Cambridge University Press, Cambridge. Runnels, C. N., and Pope, K. O. (1986), Five thousand years of land use and abuse in the southern Argolid. Hesperia 55: 103–28. Zangger, E., and Demitrack, A. (1990), Land use and soil erosion in Prehistoric and Historical Greece. Journal of Field Archaeology 17: 379–96. Vita-Finzi, C. (1969), The Mediterranean Valleys: Geological Changes in Historical Times. Cambridge University Press, Cambridge. (1976), Diachronism in Old World alluvial sequences. Nature 263: 218–19. (1986), Recent Earth Movements: An Introduction to Neotectonics. Academic Press, London. Westaway, R., Pringle, M., Yurtmen, S., Demir, T., Bridgland, D., Rowbotham, G., and Maddy, D. (2003), Pliocene and Quaternary surface uplift of western Turkey revealed by long-term river terrace sequences. Current Science 84: 1090–101. (2004), Pliocene and Quaternary regional uplift in western Turkey: the Gediz River terrace staircase and the volcanism at Kula. Tectonophysics 391: 121–69.

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Woodward, J. C. (1995), Patterns of erosion and suspended sediment yield in Mediterranean river basins, in I. D. L. Foster, A. M. Gurnell, and B. W. Webb (eds.), Sediment and Water Quality in River Catchments. John Wiley & Sons Chichester, 365–89. and Foster, I. D. L. (1997), Erosion and suspended sediment transfer in river catchments: environmental controls, processes and problems. Geography 82/4: 353–76. Lewin, J., and Macklin, M. G. (1992), Alluvial sediment sources in a glaciated catchment: the Voidomatis basin, northwest Greece. Earth Surface Processes and Landforms 16: 205–16. Macklin, M. G., and Lewin, J. (1994), Pedogenic weathering and relative age dating of Quaternary alluvial sediments in the Pindus Mountains of northwest Greece, in D. A. Robinson and R. B. G. Williams (eds.), Rock Weathering and Landform Evolution. John Wiley & Sons, Chichester, 259–83. Lewin, J., and Macklin, M. G. (1995), Glaciation, river behaviour and the Palaeolithic settlement of upland northwest Greece, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Balkema, Rotterdam, 115–29.

Hamlin, R. H. B., Macklin, M. G., Karkanas, P., and Kotjabopoulou, E. (2001), Quantitative sourcing of slackwater deposits at Boila Rockshelter: A record of Lateglacial flooding and Palaeolithic settlement in the Pindus Mountains, Northwest Greece. Geoarchaeology: An International Journal 16: 501–36. Macklin, M. G., and Smith, G. R. (2004), Pleistocene glaciation in the mountains of Greece, in J. Ehlers and P. L. Gibbard (eds.), Quaternary Glaciations—Extent and Chronology: Part 1. Elsevier, Amsterdam, 155–73. Hamlin, R. H. B., Macklin, M. G., Hughes, P. D., and Lewin, J. (2008), Glacial activity and catchment dynamics in northwest Greece: Long-term river behaviour and the slackwater sediment record for the last glacial to interglacial transition. Geomorphology 101: 44–67. Zielhofer, C., Faust, D., Baena Escudero, R., Diaz del Olmo, F., Kadereit, A., Moldenhauer, K., and Porras, A. (2004), Centennial-scale Late Pleistocene to mid-Holocene synthetic profile of the Medjerda Valley, northern Tunisia. The Holocene 14: 851–61. Zorzou, M. (2004), Suspended sediment delivery and sediment properties in mountain catchments of western Greece. Ph.D. thesis, University of Leeds.

This chapter should be cited as follows Macklin, M. G. and Woodward, J. C. (2009) River systems and environmental change, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 319–352.

12

Glacial and Periglacial Environments Philip Hughes and Jamie Woodward

Introduction Traditionally, glacial and periglacial geomorphology has not featured prominently in discussions about the physical geography of the Mediterranean basin. It is now clear, however, that on numerous occasions during the Pleistocene, and to a lesser extent during the Little Ice Age (LIA), glacial and periglacial activity was widespread in many of the region’s mountain ranges (Hughes et al. 2006a; Hughes and Woodward 2008). Even today, small glaciers and active periglacial features can be found on the highest peaks. Many mountain landscapes in the Mediterranean basin are therefore the product of glacial and periglacial processes that have fluctuated in intensity and spatial extent through the Quaternary. Glacial processes are defined here as those occurring as a result of dynamic glacier ice. The periglacial zone is sometimes defined as non-glacial areas where the mean annual temperature is less than 3◦ C (French 1996: 20). However, cryogenic processes can be important in landform development, even in areas of shallow frost over a wide range of mean annual temperatures. Thus, the term ‘periglacial’ is applied here to areas characterized by cold-climate processes—where frost and nival processes are important—but where glaciers are absent. Glacial and periglacial processes in the uplands can exert considerable influence upon geomorphological systems at lower elevations. Fluvial systems, for example, over a range of timescales have been shown to be especially sensitive to changes in sediment supply and water discharge from glaciated mountain headwaters (Gurnell and Clark 1987; Woodward et al. 2008). Nonetheless, the geomorphological impacts of glaciation

are most clearly evident in the Mediterranean mountains where the erosional and depositional legacy is frequently well preserved. Cirques, glacial lakes, icescoured valleys, moraines, pronival ramparts, relict rock glaciers, and other glacial and periglacial features can be found in many Mediterranean mountain ranges (Hughes et al. 2006a). Upland limestone terrains are widespread across the Mediterranean and many of these landscapes have been shaped by a combination of glacial and karstic processes (Chapter 10). In fact, glacio-karst is probably the dominant landscape in many mountain regions, including the Dinaric Alps of Croatia/Bosnia/Montenegro (Nicod 1968), the Cantabrian Mountains of Spain (Smart 1986) and the Pindus Mountains of Greece (Waltham 1978; Woodward et al. 2004; Hughes et al. 2006b). Glaciated mountain landscapes are also found in other rock types including granite (e.g. Monte Cinto, Corsica) (Conchon 1978) and ophiolite (e.g. Mount Smolikas, Greece) (Hughes et al.2006c). Glacial and periglacial landforms provide an important record of environmental change and can be used to reconstruct past glacial climates. This chapter examines the evidence for both past and present glacial and periglacial activity in the Mediterranean mountains and assesses its wider significance in the region. All the sites referred to in this chapter are shown in Figure 12.1 and all elevations are given in metres above sea level. This chapter is divided into three main parts. These consider, in turn, glacial and periglacial environments of the present, during the Holocene (with emphasis on the LIA), and during the cold stages of the Pleistocene.

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13 2 1

3

4 6

5

9

8

14 10

11

18

7

25 26

12

19 17

11

Currently glacierized areas

1

Other mountain areas

0

Serra da Estrela Serra do Gêres Serra de Queixa Cantabrian Mountains Sierra de Gredos Peñalara Massif

7 8 9 10 11 12

21

20

24

27

1 2 3 4 5 6

22

16 15

Sierra Nevada Pyrenees Alpes Maritimes Corsica Italian Apennines Mount Etna

13 14 15 16 17 18

23

1000 km

Julian Alps Dinaric Alps Pindus Mountains Mount Olympus Crete Pirin/Rila Mountains

19 20 21 22 23 24

Western Taurus Central Taurus Eastern Taurus Pontic Mountains Jbel Liban Aurès Massif

25 Djurdjura Massif 26 Middle Atlas 27 High Atlas

Fig. 12.1. Map of the Mediterranean showing the main mountain areas referred to in this chapter. Currently glacierized areas are indicated.

Modern Glacial and Periglacial Environments Modern glaciers are rare in the Mediterranean because most mountains do not reach the snowline. Glaciers are present, however, under modern climatic conditions in the Pyrenees, Alpes Maritimes, Italian Apennines, Julian Alps, and in the mountains of Montenegro and Turkey (Table 12.1 and Figure 12.1). These glaciers are mainly small cirque glaciers on mountains above 2,500 m, but extensive valley glaciers and small ice caps are present on some of the highest mountains (>4,000 m). Periglacial processes are much more widespread and are active above 2,000 m in areas as far south as the Moroccan Atlas and at lower elevations in the northern Mediterranean. Snowfall is heavy in many of the Mediterranean mountains with most falling during the winter months between November and March (Chapter 3).

Turkey Thirty-eight modern glaciers have been identified in the mountains of Turkey and the largest Mediterranean

TABLE 12.1. Modern glaciers in the Mediterranean Mountain area

Pyrenees Alpes Maritimes Apennines Julian Alps Turkey Montenegro

Number of glaciers 41 6 1 1 40+ 1

Total glacier area (km2 )

ELA (m)

11.43 0.31 0.045 0.0303 22.9 0.05

2,320–3,005 2,800 2,750 2,500 2,900–4,100 2,150

Maximum glacier length (km) 2.0 29,00014 C years BP3 (radiocarbon); >15,000 years4 (21 Ne cosmogenic) 21,000 years5 (36 Cl cosmogenic) Alpes Maritimes: 19,000 years6 (10 Be cosmogenic) Turkey: 26,000 years7 (10 Be cosmogenic); 20,000 years8 (36 Cl cosmogenic)

6

Late Rissian/ Saalian

Greece: >120,000 years9 (U-series) Italy: >130,000 years10 (U-series) Iberia: >130,000 years4 (21 Ne cosmogenic)

8

Early Rissian/ Saalian

Iberia: >230,000 years4 (21 Ne cosmogenic)

12

Mindelian/ Elsterian

Greece: >350,000 years9 (U-series)

1

Giraudi and Frezzotti (1997); Jalut et al. (1992). 2 García-Ruiz et al. 2003. 3 Jiménez-Sánchez and Farias (2002). 4 Fernandez Mosquera et al. (2000). 5 Palacios et al. (2007). 6 Granger et al. (2006). 7 Akçar et al. (2007), Akçar et al. (2008). 8 Sarıkaya et al. (2008). 9 Hughes et al. (2004); Hughes et al. (2006d); Woodward et al. (2004). 10 Kotarba et al. (2001).

Middle Atlas also show evidence of former glaciation and the regional snowline is estimated at c.2,800 m a.s.l. during the most extensive glacial phase (Raynal et al. 1956; Awad 1963). Periglacial features are also present. Stone polygons, solifluction features, and rock glaciers have been described on Bou Iblane and Jbel BouNaceur, in the Middle Atlas (Raynal 1952; Dresch and Raynal 1953; Awad 1963). Glacial features have also been noted in the Djurdjura Massif (2,308 m) of the Algerian Tell (Figure 12.1) where Barbier and Cailleux (1950) identified cirques, Ushaped valleys, and terminal moraines. To the southeast, in the Aurès Massif, Ballais (1983) noted the presence of moraines above 1,600 m on Jbel Ahmar Khaddou (2,017 m a.s.l.) and Jbel Mahmel (2,321 m a.s.l.). However, the chronology of glaciation in these areas and elsewhere in the Atlas Mountains has not been established and this remains the biggest obstacle to an improved understanding of the glacial history of northwest Africa.

Pleistocene Overview It is clear that very substantial ice masses formed in many Mediterranean mountain areas during Pleistocene cold stages. The glacial deposits and landforms they produced represent important archives of environmental change. However, until quite recently good dating frameworks (Table 12.3) and detailed stratigraphical frameworks had not been established in many key areas. A clear pattern is now emerging whereby the

oldest and most extensive glacial deposits and landforms date from the Middle Pleistocene. In fact, in some areas, at least two phases of Middle Pleistocene glaciation have been identified during intervals equivalent to the Saalian and Elsterian Stages of northern Europe (e.g. Fernadez Mosquera et al. 2000; Woodward et al. 2004; Hughes et al. 2006a, b). Where sound stratigraphical frameworks supported by a robust geochronology do exist for the last cold stage, it has become apparent that glacier maxima in many parts of the Mediterranean mountains preceded the global LGM of MIS 2 by more than 10,000 years (Hughes and Woodward 2008). The small mountain glaciers of the Mediterranean would have advanced and decayed rapidly in response to mass balance fluctuations and they reached their maximum during the last cold stage well before the large ice sheets that covered the Alps and northern Europe. Increased aridity in southern Europe, caused by a strengthening of high pressure systems over the expanding Alpine and north European ice sheets, would have forced glacier retreat in the Mediterranean mountains. The situation is also likely to have been complicated by millennial-scale climate changes recorded in the Greenland ice sheet (Dansgaard et al. 1993) and mirrored in long lacustrine pollen sequences in Italy and Greece (Allen et al. 1999; Tzedakis et al. 2004; Chapter 4). This has been highlighted for the mountain glaciers of Greece by Hughes et al. (2006d) who recognized the potential for multiple phases of glacier advance and retreat during the Last Glacial cycle (Figure 12.19). This analysis identified ten periods, between 115 and

Glacial and Periglacial Environments

377

Fig. 12.19. Climatically favourable conditions for glacier formation based on a combination of pollen and sea surface temperature data. Summary pollen percentage curves from the Ioannina I-284 sequence in north-west Greece, spanning the Last Glacial cycle. Potential intervals suitable for glacier formation are indicated by letters A: major stadials characterized by low arboreal pollen, both including and excluding Pinus and Juniperus; B: intermediate phases between the apices of major stadials and interstadials; and C: intervals characterized by large differences between total arboreal pollen frequencies and arboreal pollen frequencies, excluding Pinus and Juniperus. All other intervals represent major interstadials or interglacials. Both B and C types—the more favourable conditions for glaciation—are indicated by shading. The upper graph depicts variations in the percentage of Neogloboquadrina pachyderma (sinistral) and alkenone-derived seasurface temperatures in marine core MD95-2043 from the Alboran Sea, in the western Mediterranean (from Hughes et al. (2006d).

10 ka, when the climate would have favoured glacier development. The timing of the glacier maxima during earlier glaciations is unclear at present, although large glaciers in the Mediterranean may have been less responsive to rapid climate changes in comparison to those that existed during the last cold stage. Recently published data from Greece show that the transition from glacial to non-glacial conditions took place very rapidly at the end of the last cold stage (Woodward et al. 2008).

Glacial and Periglacial Interactions with Other Geomorphological Systems Glacial and periglacial systems may exert considerable influence on other environments, especially downstream fluvial systems as has been shown for Mount Tymphi, in the Pindus Mountains, where Pleistocene glaciation was a major influence on the longterm behaviour of the Voidomatis River (Chapter 11).

Fluvial sediments transported during cold stages were dominated by materials from the glaciated upland terrains (Bailey et al. 1990; Lewin et al. 1991; Woodward et al. 1992, 1995; Hamlin et al. 2000). The interaction of glacial and fluvial systems has also been explored in the Pineta basin of the central Spanish Pyrenees by Jones (2000). Glaciers and their meltwaters have enhanced karstic processes in many upland areas. In proglacial areas, dolines may be found concentrated in clusters outside terminal moraines (Waltham 1978). Where glaciers occupy only the highest valley and cirque areas, meltwaters often discharge underground before emerging as springs at lower elevations. This process is evident in front of modern and former glaciers in the Pyrenees and was termed ‘Pyrenean type’ glacio-karst by Ford and Williams (1989). This process has also been observed at the Zeleni Sneg glacier, on Triglav in Slovenia, where meltwater disappears through a 280-m deep pothole and resurfaces 1.25 km down-valley (Gams 2001). The deepest potholes in the world have formed in glaciated mountain areas such as the Pyrenees, Alps,

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and Caucasus (ibid.; Chapter 11). Karstic hollows, such as dolines, often promote snow accumulation and can contribute to glacier development. The interaction of glacial and karstic processes is an important and distinctive component of the physical geography of the Mediterranean mountains. Paraglacial effects have also been reported in the region. For example, Lebourg et al. (2003) discussed the importance of high mountain landslides and their sliding mechanisms using a case study of the glaciated Aspe Valley of the Pyrenees. They noted that slope failures occur most frequently where glacial deposits remain and generally at the junction between the till and the underlying strata (Chapter 6). Unstable landslide-prone slopes are likely to be a legacy of glaciation in many Mediterranean mountains, especially where glaciers have retreated during the Holocene and in areas of recent permafrost melting. Over longer timescales, Woodward et al. (2008) have argued that the major Middle Pleistocene glaciations in the Pindus Mountains were responsible for delivering such large volumes of sediment to the fluvial system that reworking of these materials was a major control on river behaviour throughout the Late Pleistocene and Holocene. Finally, periglacial processes also play an important role in a range of geomorphological and sedimentary systems, particularly in terms of debris supply. Frost action can deliver significant quantities of sediment to talus slopes and fluvial systems and frost shattering of rockshelter walls can be an important mechanism of coarse debris accumulation during glacial periods (Woodward 1997; Karkanas 2001 Chapter 6). Thus, even if all the remaining Mediterranean glaciers and areas of permafrost disappear in the next few decades, their legacy will remain, not just as a relict geomorphological record, but also in shaping the sediment supply dynamics in colluvial and fluvial systems.

it is possible that many glaciers will disappear in the twenty-first century. Unfortunately, in many areas, such as in Montenegro and Albania, few detailed studies have been made of recent glacier activity and further work is necessary to understand fully recent glacier dynamics across the Mediterranean (Hughes 2007). Permafrost is present in only the highest Mediterranean mountains and is especially susceptible to minor climatic variations (Gómez et al. 2001). Sporadic and discontinuous permafrost has been identified in the Sierra Nevada and Picos de Europa of Spain, the Pyrenees, the Alpes Maritimes, the Italian Apennines, the Julian Alps, and the Pontic and Taurus Mountains and periglacial processes such as nivation are present in many more mountains. These periglacial environments are also likely to diminish in the Mediterranean mountains during the twenty-first century. It is also clear that climate change will have major implications for high-mountain ecosystems in the Mediterranean region, where plants and animals adapted to cold, alpine conditions now face higher temperatures and a surge of predators and competitors (Krajick 2004; Chapter 23). Whilst contemporary glacial and periglacial environments in most Mediterranean mountains are becoming increasingly rare, the legacy of past cold conditions is widely recorded and often exceptionally well preserved. Pleistocene glacial features such as cirques, U-shaped valleys, limestone pavements, and moraines often dominate upland landscapes. These are closely associated with karstic processes in limestone uplands. A key advance of the last decade has been the wider application of radiometric dating and this has shown that glaciers were active in many areas during the Middle and Late Pleistocene. Periglacial features such as rock glaciers, scree formations, and frost-shattered bedrock are also key landscape elements in many upland areas.

Conclusions Active glacial and periglacial environments are present in several of the highest mountains in the Mediterranean and cryospheric processes been shown to have major impacts on geomorphological systems and biotic communities throughout the upland zone (above 500 m). However, it is clear that glacial and periglacial environments are becoming increasingly marginal in this region and most glaciers in the Mediterranean mountains appear to be in retreat. There is no documented evidence of sustained glacier advance in recent decades and, given current climatic trends (Chapter 3),

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Mistardis, G. (1952), Recherches glaciologiques dans les parties supérieures des Monts Oeta et Oxya Grèce Centrale. Zeitschrift für Gletscherkunde und Glazialgeologie 2: 72–9. Nicod, J. (1968), Premières recherches de morphologie karstique dans le massif du Durmitor. Meditéranée 3: 187–216. Niculescu, C. (1915), Sur les traces de glaciation dans le massif Smolica chaîne du Pinde méridional. Bulletin de la Section Scientifique de l’Academie Roumaine 3: 146–51. Ohmura, A., Kasser, P., and Funk, M. (1992), Climate at the equilibrium line of glaciers. Journal of Glaciology 38: 397–411. Palacios, D., de Andrés, N., and Luengo, E. (2003), Distribution and effectiveness of nivation in Mediterranean mountains: Peñalara (Spain). Geomorphology 54: 157–78. Marcos, J., Andres, N., and Vazquez, L. (2007), Last glacial maximum and deglaciation in central Spanish mountains. Geophysical Research Abstracts 9: 05634. Palgrave, W. G. (1872), Vestiges of the glacial period in northeastern Anatolia. Nature 5: 444–5. Pallàs, R., Rodés, Á., Braucher, R., Carcaillet, J., Ortuño, M., Bordonau, J., Bourlès, D., Vilaplana, J. M., Masana, E., and Santanach, P. (2007), Late Pleistocene and Holocene deglaciation in the Pyrenees: a critical review and new evidence from 10 Be ages, south-central Pyrenees. Quaternary Science Reviews 25: 2937–63. Palmentola, G., Boenzi, F., Mastronuzzi, G., and Tromba, F. (1990), Osservazioni sulle tracce glaciali del M. Timfi, catena del Pindo (Grecia). Geografia Fisica e Dinamica Quaternaria 13: 165–70. Baboci, K., Gruda, G., and Zito, G. (1995), A note on rock glaciers in the Albanian Alps. Permafrost and Periglacial Processes 6: 251–7. Pappalardo, M. (1999), Remarks on the present-day condition of the glaciers in the Italian Appenines. Geografia Fisica e Dinamica Quaternaria 18: 257–69. Pechoux, P. Y. (1970), Traces of glacial action in the Mountains of Central Greece. Revue de Géographie Alpine 58: 211–24. Penck, A. (1885), La Période glaciaire dans les Pyrénées. Bulletin de la Societe d’Histoire Naturelle de Toulouse 19: 105–200. (1900), Die Eiszeit auf der Balkanhalbinsel. Globus 78: 133–78. Ponel, P., Andrieu-Ponel, V., Parchoux, F., Juhasz, I., and De Beaulieu, J.-L. (2001), Late-glacial and Holocene highaltitude environmental changes in Vallée des Merveilles AlpesMaritimes, France: insect evidence. Journal of Quaternary Science 16: 795–812. Poser, J. (1957), Klimamorphologische Probleme auf Kreta. Zeitschrift für Geomorphologie 2: 113–42. Pumpelly, R. (1859), Sur quelques glaciers dans l’île de Corse. Bulletin de la Société Géologique de France 17: 78. Raynal, R. (1952), Quelques examples de l’action du froid et le neige sur les formes du relief au Maroc. Notes Marocaines 2: 14–18. Dresch, J., and Joly, F. (1956), Deux exemples régionaux de glaciation Quaternaire au Maroc: Haut Atlas Oriental, Moyen Atlas Septentrional. IV Congrès INQUA, Rome-Pisa, 1953 105–17. Ribolini, A. (1996), Note geomorfologiche sull’alta Valle del Sabbione e sulla Val’d’Ischietto (Gruppo dell’Argentera, Alpi Marittime). Geografia Fisica e Dinamica Quaternaria 19: 79–91.

(1999), Areal distribution of rock glaciers in the Argentera massif (Maritime Alps) as a tool for recent glacial evolution reconstruction. Geografia Fisica e Dinamica Quaternaria 22: 83–6. and Fabre, D. (2006), Permafrost existence in rock glaciers of the Argentera Massif, Maritime Alps, Italy. Permafrost and Periglacial Processes 17: 49–63. Robinson, D. A. and Williams, R. B. G. (1992), Sandstone weathering in the High Atlas, Morocco. Zeischrift für Geomorphologie 36: 413–29. Roth von Telegd, K. (1923), Das albanisch-montenegrinische Grenzgebiet bei Plav (Mit besonderer Berücksichtigung der Glazialspuren), in E. Nowack (ed.), Beiträge zur Geologie von Albanien. Neues Jahrbuch für Mineralogie 1. Schweizerbart, Stuttgart, 422–94. Sahsamanoglu, H. S. (1989), Mount Olympus Summer Snowfall. International Journal of Climatology 9: 309–19. Sancho, L. G., Palacios, D., De-Marcos, J., and Valladares, F. (2001), Geomorphological significance of lichen colonization in a present snow hollow: Hoya del cuchillar de las navajas, Sierra de Gredos (Spain). Catena 43: 323–40. Sarıkaya, A. M., Çiner, A., and Zreda, M. (2003), Late Quaternary glaciation of Erciyes volcano, central Turkey. XVI INQUA Congress, Reno, Nevada 23–30 July 2003, Abstracts with Programmes, Abstract 40–4: 144. Zreda, M., Çiner, A., and Zweck, C. (2008), Cold and wet Last Glacial Maximum on Mount Sandlras, SW Turkey, inferred from cosmogenic dating and glacier modelling. Quaternary Science Reviews 27: 769–80. Sawicki, R. von. (1911), Die eiszeitliche Vergletscherung des Orjen in Süddalmatien. Zeitschrift für Gletscherkunde 5: 339–50. Seret, G., Dricot, J., and Wansard, G. (1990), Evidence for an early glacial maximum in the French Vosges during the last glacial cycle. Nature 346: 453–6. Serrano, E., Agudo, C., and Pison, E. M. (1999), Rock glaciers in the Pyrenees. Permafrost and Periglacial Processes 10: 101–6. and González Trueba, J. J. (2002), La deglaciación de la laalta montaña. Morphología, evolución y fases morfogenéticas glaciares en el macizo del Posets (Pirineo Aragonés). Revista Cuaternario y Geomorfologia 16: 111–26. Serrat, D. (1979) Rock glaciers and moraine deposits in the eastern Pyrenees, in C. Schlüchter (ed.), Moraines and Varves. Balkema, Rotterdam, 93–100. and Ventura, J. (1993), Glaciers of the Pyrenees, Spain and France, in R. S. Williams and J. G. Ferrigno (eds.), Satellite Image Atlas of Glaciers of the World. United States Geological Survey Professional Paper 1386-E-2: 49–61. Sestini, A. (1933), Tracce glaciali sul Pindo epirota. Bollettino della Reale Società Geografica Italiano 10: 136–56. Sibrava, V., Bowen, D. Q., and Richmond, G. M. (eds.) (1986), Quaternary Glaciations in the Northern Hemisphere. Quaternary Science Reviews 5: 1–511. Sifrer, M. (1963) New findings about the glaciation of Triglav. Geografiski zbornik 8: 157–210. Simon, M., Garcia, I., Cabezas, O., Sanchez, S., and Gómez-Ortiz, A. (1994), Terrenos configurados ordenados en la alta montana mediterranea. Pirineos 144: 71–85. Smart, P. L. (1986), Origin and development of glacio-karst closed depressions in the Picos de Europa, Spain. Zeitschrift für Geomorphologie NS 30: 423–43.

Glacial and Periglacial Environments Smith, G. W., Nance, R. D., and Genes, A. N. (1997), Quaternary Glacial History of Mount Olympus. Geological Society of America Bulletin 109: 809–24. Smith, K. (2004), Trekking in the Atlas Mountains. Cicerone, Milnthorpe. Toro, M., Flower, R. J., Rose, N. L., and Stevenson, A. C. (1993), The sedimentary record of the recent history in a high mountain lake in central Spain. Verhandlungen Internationale Vereinigung Limnologie 25: 1108–12. Tzedakis, P. C. (1994), Vegetation change through glacialinterglacial cycles: a long pollen sequence perspective. Philosophical Transactions of the Royal Society of London B345: 403–32. Lawson, I. T., Frogley M. R., Hewitt G. M., and Preece R. C. (2002) Buffered Tree Population Changes in a Quaternary Refugium: Evolutionary Implications. Science 297: 2044–7. Frogley, M. R., Lawson, I. T., Preece, R. C., Cacho, I., and de Abreu, L. (2004), Ecological thresholds and patterns of millennial-scale climate variability: The response of vegetation in Greece during the last glacial period. Geology 32: 109–12. Vieira, G., Ferreira, A. B., Mycielska-Dowgiallo, E., Woronko, B., and Olszak, I. (2001), Thermoluminescence Dating of Fluvioglacial Sediments Serra da Estrela, Portugal. V REQUI – I CQPLI, Lisbon, 23–7 July 2001, 85–92. Mora, C., and Ramos, M. (2003), Ground temperature regimes and geomorphological implications in a Mediterranean mountain (Serra da Estrela, Portugal). Geomorphology 52: 57–72. Waltham, A. C. (1978), The Caves and Karst of Astraka, Greece. Transactions of the British Cave Research Association 5: 1–12. Woodward, J. C. (1997), Late Pleistocene rockshelter sedimentation at Megalakkos, in G. N. Bailey (ed.), Klithi: Palaeolithic

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Settlement and Quaternary Landscapes in Northwest Greece, ii. Klithi in its Local and Regional Setting. MacDonald Institute for Archaeological Research, Cambridge, 377–93. Lewin, J., and Macklin, M. G. (1992), Alluvial sediment sources in a glaciated catchment: the Voidomatis basin, northwest Greece. Earth Surface Processes and Landforms 16: 205–16. (1995), Glaciation, river behaviour and the Palaeolithic settlement of upland northwest Greece, in J. Lewin, M. G. Macklin, and J. C. Woodward (eds.), Mediterranean Quaternary River Environments. Balkema, Rotterdam, 115–29. and Smith, G. R. (2004), Pleistocene Glaciation in the Mountains of Greece, in J. Ehlers and P. L. Gibbard (eds.), Quaternary Glaciations—Extent and Chronology. Part I: Europe. Elsevier, Amsterdam, 155–73. Hamlin, R. H. B., Macklin, M. G., Hughes, P. D., and Lewin, J. (2008), Glacial activity and catchment dynamics in northwest Greece: Long-term river behaviour and the slackwater sediment record for the last glacial to interglacial transition. Geomorphology 101: 44–67. World Glacier Monitoring Service (2003), Glacier Mass Balance Bulletin 7 , accessed 23 October 2008. (2005), Glacier Mass Balance Bulletin 8 , accessed 23 October 2008. Wright, H. E. (1962), Pleistocene glaciation in Kurdistan. Eiszeitalter und Gegenwart 12: 131–64. Xoplaki, E., Maheras, P., and Luterbacher, J. (2001), Variability of climate in meridional Balkans during the periods 1675–1715 and 1780–1830 and its impact on human life. Climatic Change 48: 581–615. Yılmaz, Y., Güner, Y., and S¸ aro˘glu, F. (1998), Geology of the Quaternary volcanic centres of the east Anatolia. Journal of Volcanology and Geothermal Research 85: 173–210.

This chapter should be cited as follows Hughes, P. D. and Woodward, J. C. (2009), Glacial and periglacial environments, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 353–383.

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13

Coastal Geomorphology and Sea-Level Change Iain Stewart and Christophe Morhange

Introduction The intricate shores of the Mediterranean Sea twist and turn for some 46,000 km, with three-quarters of their convoluted length confined to only four countries— Italy, Croatia, Greece, and Turkey. Just over half the coast is rocky, much of it limestone, with the remainder encompassing almost every type of littoral environment (exceptions being coral reefs and mangrove wetlands) (Table 13.1). Such littoral diversity has long made the seaboard of southern Europe, the Levant, and North Africa a fruitful natural laboratory for studying coastal geomorphology and sea-level change. The virtually enclosed sea ensures that wave processes are generally modest and the tidal range is limited (often less than half a metre), a combination that permits observational evidence of many modern shoreline features to be related precisely to mean sea level. Consequently, relative shifts in the position of now relict coastal features can be used to track the rhythms of relative sea-level change and shoreline evolution. Such rhythms have a bearing on several aspects beyond the physical geography of the Mediterranean basin: they inform archaeological reconstructions of the past settlement and exploitation of a coastal zone that has been an important focus of human activity since Palaeolithic times; they provide testing and fine-tuning for geophysical, geodynamic, and palaeoclimatic models for the region; and they set the backdrop to contemporary societal issues, such as future sea-level rise and coastline adjustments to mass tourism, which threaten the long-term sustainability of the Mediterranean littoral. In this chapter, we review these diverse facets of

the Mediterranean coastal realm to provide a synthesis of how these shores have evolved into their present-day appearance.

Morphotectonics of the Mediterranean Seaboard The Mediterranean occupies the convergence zone between two major tectonic plates, Africa and Europe, with a third, Arabia, pressing from the east. Caught within the collisional vice of these great plates are several minor plates and crustal blocks, most notably Anatolia and Apulia. The result is a complex network of plate tectonic structures that define the general configuration of the seaboard (Figure 13.1). In particular, two major subduction systems partition the Mediterranean basin into a patchwork of minor basins and subsidiary seas (Krijgsman 2002; Chapter 1). In the eastern Mediterranean, the Hellenic arc subduction system and its former extension towards Cyprus have advanced southwards consuming the ancient Tethyan ocean floor of the Ionian and Levantine Seas, and in doing so has stretched open a major new seaway, the Aegean, in its wake. In the western Mediterranean, the Calabrian arc subduction system has similarly migrated south-eastwards destroying old Tethyan ocean floor ahead of it and rifting open a suite of successive young marine basins behind (the Alboran, Valencia, Balearic-Algerian, and Tyrrhenian) (Figure 13.1). Only the Adriatic is neither ancient ocean nor young rift, but instead is a drowned epicontinental

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TABLE 13.1. Coastal environments around the Mediterranean Sea classified into bedrock coasts and accretion coasts (which include beaches, dunes, marshes, lagoons, estuaries, and deltas)

shores of restricted seas enclosed by major land masses and island chains.

Country

Littoral Cells

Accretion (km)

%

Spain France Italy Malta Yugoslavia Albania Greece Turkey Cyprus Syria Lebanon Israel Egypt Libya Tunisia Algeria Morocco

Bedrock (km) 80 1,090 3,181 180 4,893 125 10,500 3,115 391 119 146 10 50 90 260 600 256

% 3 64 40 100 80 30 70 60 50 65 65 5 5 5 20 50 50

2,370 613 4,772 0 1,223 293 4,500 2,076 391 64 79 190 900 1,680 1,040 600 256

92 36 60 0 20 70 30 40 50 35 35 95 95 95 80 50 50

Total North coast South and east coasts

25,086 23,164

54 59

21,047 15,847

46 41

1,922

27

5,200

73

Note: More than 5% of Spain’s coastline may be described as artificial. Yugoslavia includes data for Montenegro, Croatia, and Slovenia. The north coast is defined as Spain to Turkey in the data set above (and includes all the major islands for each country) and the south and east coastline is Syria to Morocco (including Cyprus). Total length of coastline = 46,133 km. Source: Modified from Grenon and Batisse (1989).

platform. In geodynamic terms, therefore, the coastlines on and immediately inboard of the Hellenic and Calabrian arcs are the tectonically mobile borders of young active basins, while those around the periphery of the Mediterranean are, by and large, tectonically stable vestiges of the Tethyan passive margin (Figure 13.1; Chapter 1). The geodynamic complexity of the Mediterranean ensures that the three main tectonic types of coastline coexist here: collision, trailing-edge, and marginalsea coasts (Inman and Nordstrom 1971; Inman 1994; Davis 1996). Collision coasts, which characterize the narrow-shelf, mountainous seaboards of plates colliding with each other are arguably best typified by the Strait of Gibraltar collision zone where the African and European plates directly impinge. Trailing-edge coasts, which develop as wide-shelf plains on the rifted flanks of continents, characterize much of the gently warped and foundered North African margin. However, in general, the strong and pervasive tectonic partitioning in the Mediterranean means that its coastal configuration is best viewed as a nested set of marginal-sea coasts— narrow shelves fronting steep hinterlands along the

The fundamental units of study for coastal evolution, littoral cells, correspond to coastal compartments that delineate complete systems of sediment sources, transport paths, and sinks and within whose boundaries the budget of sediment is balanced (Carter 1988). The sediment dynamics of the Mediterranean littoral are strongly related to the size and character of these marginal-sea coasts. Such coasts, because they front onto smaller water bodies, are typically characterized by more limited fetch and reduced swell dynamics. In these settings, river deltas (Figure 13.2) become especially prominent and serve as important sources of sediment for littoral cells. In the southern Mediterranean, marginal-sea coasts typically exhibit large sedimentation cells, some up to hundreds of kilometres in length. Indeed, one of the world’s largest littoral cells stretches 700 km from Alexandria on the Nile delta to offshore northern Israel (Figure 13.3). Within this Nile littoral cell, sediment is swept eastwards from the delta mouth, 1 million m3 yr−1 moving by wave-dominated longshore transport and 10 million m3 yr−1 carried by coastal currents of the east Mediterranean gyre (Inman and Jenkins 1984; Chapter 2). This eastward drift of sediment is locally interrupted by eddy currents at prominent headlands such as the Damietta promontory, temporarily entraining sediment in migrating offshore sandy shoals or ribbons (Murray et al. 1981). Further east, the gradual northerly bend into the Levant coastline produces a progressive divergence in sediment flow, with the shallow longshore component becoming increasingly susceptible to wind action and feeding the extensive sand-dunefields along the coasts of the delta, Sinai, Gaza, and Israel (the ‘dry’ sink), while the shelf currents funnel the bulk into the Akhviz Submarine Canyon north of Haifa and into the Levantine basin (the ‘wet’ sink). It is no surprise that this extensive regionally simple sediment routeway has evolved along the long-lived and stable margin of the ancient Tethyan basin. Elsewhere in the Mediterranean, more active and complex geodynamics foster more complicated sedimentation cells. In the west, the Alboran Sea presents a younger and tectonically active basin in which a more restricted and intricate seaway combines with intense surface water gyres (Chapter 2) and strong winds from the Azores high-pressure cell to create smaller and more convoluted sediment transport cells

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387

(a)

Europe m

/y

r

35

mm

/yr

1 mm/yr

1m

an

uri

Lig

Ad

ria tic

cia

n Vale

BalearicAlgerian

Tyrrhenian

C Ionian

ran Albo

Aegean

H Levantine

ia

Arab

Africa (b)

Fig. 13.1. Major tectonic structures (a) and associated seismicity (b) of the Mediterranean region, highlighting how active geodynamic zones define the general coastal configuration of the region. The coastlines on and immediately inboard of the Hellenic (H) and Calabrian (C) arcs are the tectonically mobile edges of young active basins, while the coasts around the rest of the Mediterranean are generally much more tectonically stable remnants of the Tethyan passive margin (Chapter 1).

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Rhône

Po Arno Ombrone

Ebro

Adriatic

Tiber

Valencia

Axios

Basento Adra

BalearicAlgerian

Tyrrhenian

Sperkhiós

Alboran

Gediz Küçük Menderes Büyük Menderes

Achelóos

Ionian Aegean

Levantine

Sirte

Figure 13.3

Nile

Circulation pattern in the upper Mediterranean water mass Delta Characteristics

Ebro

Rhône

Po

Nile

Plio-Quaternary thickness (m) Holocene thickness Delta area (km2) Continental shelf area (km2) Shelf width (km) Shelfbreak (m) Onset of delta construction Annual discharge (tons/year) historical time Annual discharge (tons/year) modern time Delta-front advance (km/century)

1,000-1,500 30 350 9,000 65 100 early Holocene –– 3.5 × 106 5

1,000-1,500 35 720 9,000 70 100 early Holocene 60 × 106 3 × 106 4

5,000 30 770 20,000 –– 100 early Holocene –– 20 × 106 0.5-7

3,500 modified 22,000 17,000 70 250 early Holocene 140 × 106 –– 1.5

Fig. 13.2. Coastal morphodynamics of the Mediterranean basins showing the general near-surface water circulation pattern (see Chapter 2) and the locations (large triangles) and attributes (tabulated data) of the four major delta shelves: Ebro, Rhône, Po, and Nile (after Got et al. 1985). Small triangles indicate the location of other deltas discussed in the text. The box shows the area in Figure 13.3.

(Goy et al. 2003). Most of the southern coast of Spain trends east–west, roughly parallel to the prevailing winds, and this situation favours longshore currents and littoral drift. However, where prominent bays and headlands break this trend, the sediment transport routes are intercepted and occasionally locally reversed, thereby delimiting second- and third-order littoral cells. Within the even more tectonically fragmented coastal landscape of the circum-Aegean Sea (Figure 13.1), faultbounded gulfs and uplands produce a complex shelf and nearshore bathymetry that creates variable patterns of sediment routing over distances of a few tens of kilometres to kilometres. These often switch in character across active tectonic structures creating a heterogeneity of sediment sources and sinks and a collage of littoral environments (Leeder et al. 1991; Collier et al. 1995).

Tectonics, Climate, and Sea Level As well as partitioning littoral sedimentation systems, geodynamic processes determine whether the Mediterranean shores are emerging or subsiding (Milliman 1992). The fastest geodynamic movements are occurring at rates of 30–40 mm per year horizontally and about 1 or 2 mm per year vertically, which mean that over thousands to millions of years their cumulative effects make dramatic changes to the coastal configuration. The most pervasive effect of this has been the gradual closure of the seaway by Africa– Europe convergence, a process responsible for triggering a regional dessication event, the Messinian Salinity Crisis (Hsü et al. 1973), which in turn predicated the large-scale reshaping of the Mediterranean seaboard (Chapter 1).

Coastal Geomorphology and Sea-Level Change

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Fig. 13.3. The Nile littoral cell extends along the south-eastern Mediterranean coast from Alexandria, Egypt, to the Akhziv Submarine Canyon, Israel. Sediment transport paths are shown by dark arrows (modified from Inman and Jenkins 1984).

Several million years of subduction-related tectonic emergence of the Betic-Rif area had gradually shallowed the narrow seaways that connected the Mediterranean to the Atlantic Ocean (Krijgsman et al. 1999; Warny et al. 2003; Duggen et al. 2003, 2004). Around 5.96 million years ago, this tectonic uplift, aided perhaps by global glacio-eustatic processes (Clauzon et al. 1996; Hodell et al. 2001), finally closed the Atlantic gateway. Over the course of the next few hundred thousand years, evaporation of the now largely landlocked basin dropped water levels by at least 800 m, and probably as much as 1,300 m, below their modern equivalent (Ben Gai et al. 2005). Thick sequences of evaporites were deposited in the hypersaline abyssal plains. Around the shores of the Mediterranean salt lake the plummeting base-level triggered a major phase of river downcutting, creating vast planation surfaces and carving extensive

canyon networks. The most formidable river, the Nile, cut a canyon that was three times longer than the Grand Canyon with a similar depth and width (Said 1981), and major rock-cut gorges and cataracts are known to exist beneath the Rhône, the Ebro, and the Po rivers. Pronounced coastal progradation at these canyon mouths would form the earliest submarine cones of the region’s great deltas. Today, infilled palaeo-valleys can be found along much of the Mediterranean coast and its continental shelf, vestiges of this Messinian incision (Chapter 1). This ‘great drying’ ended about 5.3 million years ago, when the marine gateway to the Atlantic was restored, due to the westward propagation of the Alboran rift basin (Duggen et al. 2003, 2004) and/or eastward piracy of the Atlantic waters (Blanc 2002; Loget et al. 2005) breaching the Betic-Rifian land bridge at Gibraltar to create the present-day straits.

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Since the Messinian Salinity Crisis, the Mediterranean waters have been maintained through a fragile balance between tectonics and climate, much of it dependent on east–west differences in salinity that reflect a water budget deficit in which outputs from evaporation exceed inputs from precipitation and freshwater influx (Chapters 2 and 8). The long-term effects of this in terms of sea-level changes during Plio-Pleistocene times are poorly constrained, because the Mediterranean lacks both the sensitive marine-continental switches of the neighbouring Red Sea (e.g. Siddall et al. 2004) and the coral-reef staircases, which in locations such as Barbados, the Huon Peninsula, and Tahiti, have yielded long-term eustatic records (for a review, see Lambeck and Chappell 2001). Nevertheless, some constraints on Pleistocene sea-level oscillations have been provided by stratigraphical sequences in subsiding coastal plains and shallow shelves (e.g. van Andel et al. 1990) or in drowned littoral caves (Fornós et al. 2002; Tuccimei et al. 2003; Antonioli et al. 2004). On emerging coastlines, records of sea-level highstands have been derived from flights of marine terraces (e.g. Keraudren and Sorel 1987; Goy and Zazo 1988; Dumas et al. 1993; Carobene and Dai Pra 2003; Zazo et al. 1999, 2003; Rodríguez-Vidal et al. 2004; Ulu˘g et al. 2005).

Coastal Tectonics As well as revealing long-term (104 to 106 year) sealevel records, the ancient shorelines of the Mediterranean can also be used as a measure of geodynamic activity. A particularly useful marker is the shoreline that formed during the Last Interglacial period, 120,000–130,000 years ago, the climatic optimum of Marine Isotope Stage (MIS) 5.5. This episode has left an especially indelible trace in the Mediterranean, in part because global sea levels stood 3–6 m higher than present, and in part because the unusually warm climatic conditions favoured the development of distinctive faunal assemblages. The fossil Strombus bubonius is an index marker of this highstand (Bordoni and Valensise 1998). Figure 13.4a shows the elevation of this geomorphic marker across the Mediterranean realm. In the Strait of Gibraltar, direct continental collision has locally raised the MIS 5.5 terrace to 20 m above sea level, but this drops abruptly back to being within a few metres of modern sea level along most of Spain’s Atlantic and Mediterranean coasts (Zazo et al. 1999, 2003). Other plate-boundary structures coincide with elevated MIS 5.5 terrace elevations, notably where the horizontally slipping North Anatolian Fault intersects the Mar-

mara Sea coastline (Yaltirak et al. 2002), and more significantly in north-eastern Sicily and south-western Calabria (Miyauchi et al. 1994; Bordoni and Valensise 1998; Ferranti et al. 2006), which appear to be on the rise following the Late Pleistocene detachment of the subducted African slab (Westaway 1993). Elsewhere along the Africa–Europe subduction and collisional front, however, MIS 5.5 shorelines are roughly at their original elevation. Perhaps surprisingly, the highest Last Interglacial terraces in the Mediterranean fall not on a plate-bounding but on a prominent interplate rift zone, the Gulf of Corinth fault system (Keraudren and Sorel 1987; Armijo et al. 1996). No other rifted gulfs of the circum-Aegean coast attain such high Last Interglacial shoreline elevations (Kelletat et al. 1976), and along the opposing Turkish Aegean coast no MIS 5.5 terrace sites have so far been reported (Brückner et al. 2004). Along the Mediterranean coast of Turkey, Pirazzoli (1991) attribute higher shorelines to MIS 5.5 but these have not been dated. The general tendency for Last Interglacial marine terraces to lie within a few metres of present-day sea level indicates general long-term land stability throughout most of the circum-Mediterranean region. The marked variability in terrace elevation highlights the role of local (basin-bounding) rather than regional geodynamics. Both these characteristics are also apparent in the Holocene and modern tide-gauge records from the Mediterranean (Flemming 1978, 1993; Emery et al. 1988; Zerbini et al. 1996). According to these recent sea-level data, most of the Mediterranean coast appears to be submergent, although mostly at a low rate (−1.2 mm/yr, Emery et al. 1988). The zones of little or no sea-level change correspond to the coastal tracts between Gibraltar and Genoa in the west and along southern Turkey in the east. In contrast, the most mobile shores are mostly on or immediately inboard of the Hellenic and Calabrian arc areas of the central northern Mediterranean, where they coincide with zones of high earthquake activity (e.g. Pirazzoli et al. 1996a; Stewart et al. 1997; Chapter 16) or active volcanic centres (Dvorak and Mastrolorenzo 1991; Firth et al. 1996; Stiros 2000; Morhange et al. 2006; Chapter 15). The overall picture, therefore, is of considerable variation in both the amount and sense of vertical coastal movements along the Mediterranean littoral.

Isostatic Responses to Ice-Ocean Loading In addition to tectonic deformation, land movements in the Mediterranean littoral are subject to

Fig. 13.4. (a) Elevation of the Last Interglacial (Marine Isotope Substage 5.5) shoreline based on a compilation of published data. Terrace data from Antonioli et al. (2006), Conchon (1975), Kelletat et al. (1976), Bordoni and Valensise (1998), Zazo et al. (1999, 2003) and Yaltirak et al. (2002). Solid lines trace the main geodynamic arcs (Ca = Calabrian, Cy = Cyprus, G = Gibraltar, H = Hellenic), with solid triangles indicating active subduction fronts and open triangles denoting collisional fronts. (b) Rates of late Holocene crustal movement derived from sea-level curves in Pirazzoli (1991) augmented by more recent sea-level studies. Black downward arrows denote subsiding areas and grey upward arrows indicate emerging areas. (c) Predictions of global isostatic adjustment made for Mediterranean tidegauge stations, updated from predictions of Peltier (2001) by adopting the new analysis of Peltier (2004) and using values listed at <www.pol.ac.uk/psmsl/peltier/index.html>, accessed 27 October 2008.

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glacio-hydro-isostatic crustal effects, whereby the marine basin adjusts to past and ongoing fluctuations in surface loads of ice and water. Various numerical models have emerged which make competing predictions on the shoreline responses to ice-ocean loading (Lambeck 1995; Peltier 1998, 2000; Lambeck and Bard 2000). The intricacies of these ice-history/earthrheology models are outside the scope of this chapter, but there are essentially two key contributions to land movement. The first, a glacio-isostatic contribution, comes from rebound of the former ice-mass centres of Europe and North America, whereby mantle material squeezed beneath the Mediterranean crust by ice-sheet depression now flows back causing the previously upwarped Mediterranean ‘forebulge’ to subside. Figure 13.4c shows the predicted effect of this glacial isostatic adjustment in the Mediterranean. The second, a hydro-isostatic contribution, results from meltwater from ice-sheet decay increasing the water load of the global oceans and seas, thereby downwarping the marine basin floors and upwarping their margins (Lambeck et al. 2004a, 2004b). The application of these models to the Mediterranean littoral is discussed in Pirazzoli (2005), but key points are

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noted here. In some models, the combined effect of both glacio-isostatic and hydro-isostatic changes is to accentuate general subsidence across much of the central Mediterranean and induce compensatory upwarping of its margins (Lambeck 1995; Lambeck and Purcell 2005). However, according to Peltier (1987) the Mediterranean region is sufficiently remote from the North European centre of glaciation as to be little influenced by the collapse of its forebulge; instead, relative sea-level movements ought to be more significantly influenced by the local water load emplaced on the basin itself. The models also highlight local idiosyncracies. For instance, along parts of the North African and Levantine coast, Lambeck and Purcell (2005) predict that the two isostatic contributions can potentially counteract each other, whilst in the northern Adriatic the Alpine deglaciation probably amplifies the glacial isostatic signal (Lambeck and Purcell 2005). In both areas, accentuated coastal emergence is postulated, making it possible that mid- to late Holocene land uplift may have temporarily outpaced sea-level rise to create locally a highstand shoreline. Figure 13.5 shows sea-level and shoreline predictions for the Mediterranean region for four time slices at 20 ka, 12 ka, 6 ka, and 2 ka

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Fig. 13.5. Predicted relative sea levels and shorelines across the Mediterranean region at four epochs: a = 20 ka, b = 12 ka, c = 6 ka, and d = 2 ka. The palaeoshoreline positions are defined by the seaward side of the grey shading. For 20 and 12 ka BP, the contour intervals are 5 m. For 6 and 2 ka BP the solid contours denote negative values, the finely dashed contours denote positive values, and the dashed contours correspond to zero change (modified from Lambeck and Purcell 2005).

Coastal Geomorphology and Sea-Level Change

respectively. Note the expansion of the coastal plain at the global LGM when sea level was around 120 m lower than today. In general, however, glacio-isostatic model simulations confirm the findings of most sea-level studies, i.e. that postglacial Mediterranean sea levels have never been higher than at the present day. In other words, any signs of Holocene coastal phenomena found elevated above the modern waterline is strong evidence for local tectonic movements. It is a reminder that geological evidence for past sea levels provide the testing ground for these global isostatic models (e.g. Pirazzoli 1998). Thus, in a comparison between several late Holocene sea-level histories and glacio-hydro-isostatic predictions, Pirazzoli (2005) contends that, in non-tectonic areas, hydro-isostatic effects have been, in places, overestimated (e.g. Sardinia) and in other places underestimated (e.g. southern Tunisia). Pirazzoli (2005) disputes predictions of isostatic subsidence along the Hellenic arc and on the Levant coast. Resolving the discrepancies between model and field data will continue to refine our understanding of Mediterranean postglacial shoreline change, and in the following sections, we discuss how such geologically derived sea-level constraints are determined.

Postglacial Sea-Level Changes Tectonic and glacio-hydro-isostatic effects ensure that there is considerable spatial and temporal variability in land movements within the Mediterranean realm, but the general pattern of postglacial eustatic sea-level rise is adequately known (Lambeck and Chappell 2001). This is known thanks mainly to ‘global’ sea-level curves based upon coral-reef reference sites (e.g. Barbados, the Huon Peninsula, and Tahiti) which show that world sea levels rose by 120–30 m, largely between 16,000 and 8,000 years BP, thereafter stabilizing at present levels by 6,000 years BP (Pirazzoli 1991). However, tracking sea-level behaviour following this mid-Holocene stabilization is more difficult, because corals are low-precision reference markers whose vertical range of repartition in modern reefs is of an equivalent magnitude to eustatic variations over the last 6,000 years (Blanchon and Shaw 1995). This mid- to late Holocene period, however, is critical for isolating various non-eustatic dynamic factors such as isostasy, tectonics, and sea-surface topography that may induce significant local sea-level fluctuations (e.g. Mörner 1996). The Mediterranean littoral, however, offers an opportunity to establish such detailed late Holocene sea-level histories because it combines a diversity of potentially high-

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precision sea-level proxies with a microtidal regime. As is discussed below, the evidence of former shorelines around the Mediterranean region is key to establishing the local relative sea-level trajectories needed to refine tectonic and glacio-isostatic models.

Evidence of Former Sea Levels in the Mediterranean Littoral Zone A wide range of Mediterranean shoreline indicators can be employed for establishing Holocene sea-level curves ranging from precise ‘sea-level index points’ that securely tie the elevation of the land–sea interface at a specific period in time, to more equivocal and/or less well-dated indicators of former submergence or emergence (Haslett 2000). General reviews are provided by van de Plaasche (1986) and Pirazzoli (1991), but in brief they include a suite of sedimentological, geomorphological, geoarchaeological, and biological proxies.

Sedimentological Proxies Sedimentological evidence provides the most widespread but typically least precise record of shifting shorelines. Fossiliferous sands containing gastropods and bivalves equivalent to those of the modern infralittoral zone can be used, though altitudinal errors may be as large as 10 m. Such metre-scale precision can also be expected from cores from coastal sediment sequences using grain size or fossil fauna to identify contrasting littoral environments (such as terrestrial floodplains, freshwater lakes, seasonal pools, brackish lagoons, beaches, marine delta fronts, and prodelta slopes) and assess the extent of altitudinal change. Even where sharp interfaces between intertidal muds and subaerial peats provide reliable sealevel index points, problems of sediment compaction generally maintain altitudinal uncertainty. In Provence, for example, Vella and Provansal (2000) analysed basal peat formations from the eastern limit of the Rhône delta above a non-deformable substrate and determined an altimetric precision of ±50 cm at best. Clastic shoreline deposits (beachrock) have been used as sea-level indicators where they contain early diagenetic carbonate cements whose textures can be diagnostic of intertidal cementation; though an error of +1 and −5 m may be expected (Fouache et al. 2005b). A potentially better positional accuracy can be obtained from the drowning of littoral caves, whereby speleothems precipitated in emergent caves during lowstands subsequently become encrusted by marine biogenic overgrowths when flooded with sea water (Vesica et al. 2000;

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Fig. 13.6. Schematic representation of a littoral karst cave based on observations on Mallorca Island and Capo Caccia (Sardinia). Note the presence of phreatic overgrowths on speleothems related to former and present-day sea levels (modified from Tuccimei et al. 2003).

Bard et al. 2002; Antonioli et al. 2004; Chapter 10) (Figure 13.6). Complexities include the geometry of the cave system and the fact that there is often a hiatus of several millennia between the end of speleothem formation and marine overprinting caused by the influx of fresh groundwater ahead of saline inundation (Suri´c et al. 2005).

Geomorphological Proxies The geomorphological form of the coast reflects the local balance of bioerosion and bioconstruction. On rocky shores, bioerosion at mean sea-level forms midlittoral notches and tidal platforms. These are best developed on limestone, where they are carved by boring and grazing organisms and dissolved and abraded by wave action. Being a combination of physical, chemical and biological erosion, the nature of the local context strongly controls their precise form (Pirazzoli 1986a; Kershaw and Guo 2001; Naylor and Viles 2002; Trenhaile 2002). The effects of variations of coastal wave energy, for example, produce a continuum of notch forms whose precise relationship to mean sea-level can, in some instances, be difficult to assess (e.g. Kershaw and Guo 2001) (Figure 13.7). Nevertheless, notches and associated platforms are well developed around many carbonate coasts in the Mediterranean, and whether

emergent (e.g. Pirazzoli et al. 1994a, 1996a; Stewart et al. 1997; Rust and Kershaw 2000) or submergent (Faivre and Fouache 2003; Antonioli et al. 2006), they are valuable markers of land movements relative to sea level.

Geoarchaeological Proxies The long human legacy of the Mediterranean means that cultural features are some of the most reliable shoreline markers for reconstructing sea-level changes over recent millenia. Indeed, arguably the best example of the postglacial sea-level highstand comes from Cosquer Cave, southern France, where Palaeolithic wall paintings depicting horses have partially been eroded by the rising sea water, clearly showing that Holocene sea level has never reached higher than its present-day level (Morhange et al. 2001; Figure 13.8). Elsewhere it is archaeological remains that have long provided the reference datum (Flemming 1969, 1978 1993; Blackman 2005): for example, the study of ancient Mediterranean harbours has emerged as an important component of geoarchaeology (Marriner and Morhange 2007). Ancient rock-cut installations such as quarry floors and piscinae (fish ponds) show that the Israeli coast has been tectonically stable for the last two thousand years

Coastal Geomorphology and Sea-Level Change

(a)

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Wave amplitude

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High tide Mean sea level Low tide

Very sheltered sites

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(b)

Fig. 13.7. (a) Effects of variations of coastal-wave energy on marine-notch formation (modified from Rust and Kershaw 2000 and Pirazolli 1986a). 1 and 2 are wave notches formed in quiet conditions at mean sea level; 3 and 4 are surf notches formed in more turbulent conditions as much as 2 m above sea level. (b) A marine notch at Capo Milazzo in north-eastern Sicily illustrating the contrasting morphology of an open (right) to a sheltered (left) position (photo: Iain Stewart).

(Galili and Sharvit 1998; Sivan et al. 2001). Similarly, Lambeck et al. (2004b) have utilized Roman piscinae constructed on or in bedrock along the Tyrrhenian coast of Italy to determine that sea level here two millennia ago was 1–2 m lower than today, a change that they attribute almost entirely to glacio-isostatic subsidence. This result is out of step with previous studies (e.g. Pirazzoli 1976).

Biological Proxies

Present mean sea level

Sublittoral zone

Fig. 13.8. The absence of a Holocene sea level above present datum is supported by the evidence of painted horses on a wall of a half-submerged Palaeolithic cave near Marseilles. Below the current water level of Cosquer Cave the wall paintings are significantly degraded (after Morhange et al. 2001).

The above proxy indicators are by no means independent—ancient harbours, for example, are important stratigraphical archives (e.g. Kraft et al. 2003; Marriner et al. 2006; Marriner et al. 2006)—and many Mediterranean sea-level studies typically combine sedimentological, geomorphological, and archaeological indicators. However, few of these indicators are valuable

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without associated biological indicators. As well as providing often precise reference markers for palaeo-sealevels, biological proxies also serve as dateable deposits from which to establish sea-level histories. Over the last decade or so, the use of biological sea-level indicators in the study of Mediterranean sea-level changes has gradually evolved from a descriptive to a multidisciplinary approach integrating many of the proxies above (Laborel and Laborel-Deguen 1994, 1996). It is an approach based on the recognition that the vertical distribution of the littoral fauna and flora of rocky shores shows a pattern of superimposed ecological belts, a tendency called biological zonation (Péres 1982). According to biological zonation, marine benthic animals and plants are finely adapted to very precise ecological conditions such as light intensity, turbidity, water salinity and temperature, and surf exposure. Consequently, changes in local ecological conditions are followed by a concomitant quantitative and qualitative modification of the organisms with replacement by more tolerant forms. Laborel (1986) discusses this biological zonation in detail and shows how it can be used to measure past sea levels. In very general terms, several parallel zones can be recognized (Figure 13.9) and these are outlined as follows: 1. A supralittoral fringe (or supralittoral zone) never or rarely submerged but wetted by surf in which the biomass is very low and mainly represented by boring endolithic cyanobacteria and grazing gastropods. 2. A midlittoral zone submerged by tides and waves on a regular basis, which displays a pattern of

(a)

parallel algal and faunal belts, with biomass and species diversity increasing downwards. Cyanobacteria, limpets (Patella spp.) and Chitons are the main bio-eroders in this zone. Constructional elements such as the rim-building coralline rhodophyte Lithophyllum byssoides may develop in the north-west Mediterranean. 3. An infralittoral (or sublittoral) zone whose upper limit is marked by a sudden increase in biodiversity (Boudouresque 1971), thus defining a biological sea level that ranges down to the lower limit of marine phanerogams and photophilous algae, i.e. to a mean depth of about 25–35 m. The upper part of this infralittoral zone (also called ‘infralittoral fringe’) is densely populated by brown algae (Cystoseira), Coralline Rhodophytes, fixed vermetid gastropod molluscs (such as Dendropoma sp.), and cirrhipeds, for example Balanus spp. Active erosive agents, such as clionid boring sponges, sea-urchins, and rock-boring mussels (Lithophaga, Hyatella, Coralliophaga spp.), are responsible for rapid underwater erosion of the limestone outcrop. The limit between the midlittoral and the infralittoral corresponds to the ‘biological sea level’(Laborel 1986). The influence of local variations in coastal morphology upon surf exposure explains why this limit may be slightly undulating locally. Aperiodic sea-level oscillations linked to atmospheric pressure or wind variations have little influence upon the marine zonation of living organisms with a lifespan of more than one year. Biological sea level itself is best characterized by

(b) Action

Agent

Erosion construction

Rain water Sea spray Chfthamalus Cyanobacteria, limpets Lythophyllum byssoides Protectio rom Protection from brown algae Sea urchins Clionas Erosion Lithophaga

Resulting morphology

Biological zones

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Biokarst

Biological zones

Resulting morphology

Supralittoral

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Upper midlittoral Mid midlittoral

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Notch

Lower midlittoral

Lower midlittoral

Platform

Subtidal

Biokarst

Subtidal

Erosion Agent construction

Action

Rain water Sea spray

Biokarst

Chfthamalus Cyanobacteria, limpets Dendropoma petraeum Vermetids Sea urchins Clionas

Boring Molluscs Fig. 13.9. A schematic coastal profile showing the main characteristics of bioconstruction and biodestruction on calcareous coasts in (a) the western Mediterranean, and (b) the eastern Mediterranean (modified from Laborel and Laborel-Deguen 1994).

Coastal Geomorphology and Sea-Level Change

the development of the few marine species, with a very narrow depth range, located immediately above (e.g. Lithophyllum byssoides rim) or below (e.g. Dendropoma rim) which are considered the more precise sea-level indicators. When such species are absent, the highest limit of biological perforations by Cliona and Lithophaga (Laborel and Laborel-Deguen 1994) are also excellent proxies. Consequently, littoral algal and vermetid bioconstructions and the upper limits of bioerosive elements (marine burrows and perforations), and fixed invertebrates (oysters, barnacles, solitary vermetids) are commonly used as biological sea-level indicators.

Biological Evidence for the Amplitude and Rapidity of Shoreline Change If a suitable biological indicator is available, and if an accurate study of local ecological conditions has been undertaken, fossil biological sea-level markers can provide reliable constraints on the direction, amplitude, rapidity, complexity, and timing of apparent coastal deformation (ibid.). Of these, indicators of the amplitude and rapidity of relative sea-level displacement are particularly important for discerning tectonic versus eustatic sea-level changes. The amplitude of shoreline change is derived from the altitude of elevated or submerged shorelines. This parameter is best estimated by direct measurement of the altitudinal difference between the upper limit of the elevated (or submerged) remains and the corresponding upper limit of its modern equivalent. This can easily be done with species such as Dendropoma sp. or Lithophyllum byssoides whose populations have a very narrow vertical range closely connected to biological sea level. For species with a wide vertical range such as Lithophaga, balanids, or solitary vermetids, good results are obtained only when the uppermost limits of both fossil and living populations are well delineated or when the fossil remains are correlated with a morphological sea-level indicator such as a tidal notch. Any direct reference to the present instantaneous water level, whether observed or calculated from tide tables, is thus unnecessary and incorrect. For example, the only Mediterranean reefcoral, Cladocora caespitose, can grow from the surface down to more than 40 m, but it does not present any well-marked upper limit nor does it develop distinct sealevel linked build-ups. It should therefore only be used as a proof of submersion but not as a depth proxy and should not be attributed with a precise depth coefficient (Antonioli et al. 1999). When a series of measurements is to be made in a limited area, one must keep in mind as stated above,

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that a biological sea-level mark is not a perfect horizontal line but may be naturally warped, even over very short distances, by local hydrodynamic variations. For that reason, each individual altitudinal measurement must be done on a vertical profile of its own, including both the fossil specimen and its present equivalent (Laborel 1979). The accuracy that can be obtained depends upon both the state of preservation of the upper limit of fossil populations and local ecological conditions. An accuracy of about ±5 cm was obtained in south-west Crete for a series of remarkably well-preserved vermetid rims (Thommeret et al. 1981). In Euboea, for elevated populations of Lithophaga burrows (Stiros et al. 1992), the lowest accuracies observed (around ±50 cm to several metres) are typically found in relation to Chthamalus populations. Biological sea-level indicators are especially useful in determining the rapidity of relative sea-level changes. In the case of slow uplift (less than a few millimetres per year), biological indicators living in the subtidal zone are killed by emersion and their remains are slowly carried up through the midlittoral zone to the supralittoral zone. Small or fragile skeletons quickly disappear by bioerosion and physical abrasion. Massive algal rims and vermetid constructions are not completely destroyed and can be used as proxies. In contrast, good preservation of fragile elevated remains is the best evidence of a rapid uplift. Determining the rapidity of subsidence is much more difficult than for uplift, because many biological sea-level indicators are either rapidly destroyed by subtidal erosion or covered by new generations of marine organisms. Thus, only bioconstructing species with a very narrow vertical range, such as Lithophyllum byssoides or Dendropoma petraeum., may be of some use. For such bioconstructed rims, it is often possible to obtain valuable information from the distribution of their drowned remnants, provided they are not very old and have not been removed by biological erosion. Eroded remains below the present rim are often found where subsidence is slow. More rapid change would not provide sufficient time for the development of a bioconstructed rim at intermediate depth, and drowned bioconstructions would thus not appear clearly separated (Laborel and LaborelDeguen 1994).

Holocene Sea-Level Histories for Contrasting Mediterranean Coasts As discussed earlier, there is considerable variation in the stability of the Mediterranean coastline, with the

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Fig. 13.10. Age-depth diagram from Marseilles’s archaeological excavations compared with dated algal rims from nearby rocky cliffs (modified from Morhange et al. 2001 and Laborel and Laborel-Deguen 1994). NGF = Nivellement Général de la France.

geodynamically active Calabrian and Aegean seaboards being especially mobile, the rifted margins of the southern, western, and eastern sectors being largely stable, and localized volcanic centres experiencing irregular paroxysms (Figure 13.4b). In the following section, we examine in more detail the relative sea-level histories established for three sites in these contrasting settings.

‘Stable’ Coast The site of Marseilles (southern France) is a good example of relative sea-level ‘stabilization’ in a socalled tectonically ‘stable’ area (Morhange et al. 2001). Allied with recent archaeological excavations of the ancient harbour, biological indicators have yielded highprecision data for the past 5,000 years (Figure 13.10). One of the best biological sea-level indicators used here is the upper limit of barnacles (Balanus sp.). They commonly develop upon quay walls in clear or polluted waters, stopping abruptly at biological sea level. When their upper limit is continuous, a precision of plus or minus a few centimetres can be obtained. Wherever barnacle-bearing hard surfaces were not available, the upper-limit of subtidal beach onlap layers was used as a sea-level indicator with a precision of ±0.2 m.

Data obtained in Marseilles fit well with indicators from rocky coasts such as Lithophyllum rims (Laborel and Laborel-Deguen 1994). The age-depth diagram shows a regular sea-level rise up to about AD 500 followed by a period of stability. Total rise has been less than 1.5 m since 4,500 years cal. BP. Observations do not show any Holocene level higher than present (Figure 13.10). The rate of mean sea-level rise was 0.4 mm/yr between 4,500 years cal. BP and AD 500, and 0.2 mm/yr thereafter. During the twentieth century, the rate of sea-level rise increased to c.1.5 mm/yr, most likely in connection with global warming. In other words, in settings where geodynamic activity is low or insignificant, background eustatic trends can be isolated.

Volcanic Coasts More complicated coastal elevation changes can occur in areas affected by volcano-tectonic deformation (e.g. Firth et al. 1996). The most detailed record of the complex mobility of volcanically active shores can be shown by the archaeological ruins of Pozzuoli on the Phlegrean Fields in the Bay of Naples. The bioeroded columns of the port’s Roman market were made famous as the frontispiece of Charles Lyell’s Principles of Geology, and

Coastal Geomorphology and Sea-Level Change

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Estimated curve (biological indicators) Biological zero 1995

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Fig. 13.11. Measured relative sea-level changes in the old harbour of Pozzuoli compared to estimated relative sea-level changes using biological indicators (modified from Morhange et al. 2006).

since that time have served as a palaeo-tide gauge to track the deformation history for the area (Dvorak and Mastrolorenzo 1991; Orsi et al. 1999; Morhange et al. 2006). Relative sea-level movements at Pozzuoli are far more intense than the average 50-cm sea-level rise recorded in the north-west Mediterranean sea since Roman times (Pirazzoli 1976). Moreover, the post-Roman sea-level history reveals three cycles of submersion and emergence (Figure 13.11). The first submersion accompanied a period of marine transgression which ended around AD 400–550 without any volcanic eruption. A second submersion affected the site in the early Middle Ages, around AD 700–900, though again no eruptive activity followed this inundation. The third submersion occurred in the late Middle Ages, when it was followed by a well-documented pulse of land uplift that culminated in a major eruption in 1538. More recently, the two major events of 1969–72 and 1982–4 resulted in a total uplift of c.3.5 m, although no eruption occurred during these periods. Indeed, during the last 2,000 years, noneruptive coastal uplift episodes have been the rule rather than the exception. Clearly, here, the dominant sea-level change signal is linked to phases of volcanic activity and quiescence (see Chapter 15).

Pirazzoli et al. 1994a, 1994b, 1996a; Stewart 1996; Stewart et al. 1997). Such ‘coastal palaeoseismology’ studies reveal that seismically active coasts in the Mediterranean realm typically experience decimetreto metre-scale seismic jerks every few centuries to millennia. Arguably the most remarkable example of this seismotectonic action has been reported from the shores of western Crete and nearby Antikythira island, where a series of elevated stepped shorelines are carved into the limestone cliffs supporting well-preserved Vermetid formations. The highest of these contains very fragile shell skeletons that would have been rapidly destroyed in the surf zone (Thommeret et al. 1981). The sea-level history derived from these shorelines is especially complex (Figure 13.12). It reveals that, between 4,000 and 1,700 years BP, ten successive increments of small but abrupt subsidence affected a huge block of lithosphere about 150 km long without noticeable tilting. This phase of subsidence was followed about 1,530 years BP by an abrupt uplift, by about 9 m, of the south-west corner of Crete and a north-eastward tilting of the western

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Tectonically Active Coast In contrast to volcanically active coastlines where land movements are discernible over months to years, equivalent shoreline movements on earthquake-prone coasts tend to be instantaneous. Abrupt littoral emergence or submergence in these environments has been widely used to identify past large earthquakes in the Holocene coastal record (e.g. Stiros et al. 1992;

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Fig. 13.12. Recent relative sea-level variations in Antikythira island (Greece), from Pirazzoli (1986b) and Pirazzoli et al. (1996b).

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part of the island. In north-west Crete, the Roman harbour of Phalasarna was raised by over 6 m (Pirazzoli et al. 1992; Chapter 16). Relative sea-level change histories at other coastal sites from western Greece to the Levant coasts also appear to exhibit marked vertical tectonic displacements—some of them several metres in amplitude—between 1,750 and 2,000 years BP (Pirazzoli 1986b; Pirazzoli et al. 1996b). These studies attribute the main movement to a great earthquake in AD 365 (Stiros and Papageorgiou 2001), though recalibration of the original radiocarbon dates on elevated littoral fauna led Price et al. (2002) to shift the main emergence event to the 6th century AD, another period of seismic unrest on Crete (Di Vita 1996). The tectonic mechanism for such dramatic coastal displacements also remains uncertain, but they have been ascribed to an exceptional burst of tectonism that occurred on a regional scale in post-Roman times—the

‘Early Byzantine Tectonic Paroxsym’ (Pirazzoli 1986b; Chapter 16). Nevertheless, the literally Byzantine history of Holocene shoreline development in western Crete highlights the potential complexity of land movements at plate boundary zones.

Tsunamis as a Coastal Process The high seismicity of the Mediterranean region means that much of its coastline is subject to recurrent tsunamis (Soloviev 1990; Chapter 17). Unsurprisingly, the main areas of observed tsunami incidence coincide with the principal earthquake belts (Figure 13.13). However, in addition, volcanic-induced tsunamis affect the Tyrrhenian and Aegean seas, and the steep margins of the north-western Mediterranean basin mean that tsunamigenic submarine slumps are potentially commonplace. The result is that while some coastal zones are more susceptible than others, virtually any

Fig. 13.13. Tsunami activity in the Mediterrenean Sea. Historical data compiled from the following sources: Ambraseys (1962) for the Levant; Soloviev (1990) for the western Mediterranean; Tinti and Maramai (1996) for the French Côte d’Azur and Italy; Papadopoulos (1998) for Greece and the Aegean; Altnok and Ersoy (2000) and Yalçıner et al. (2002) for Turkey. Palaeo-tsunami sites are compiled from Wood (1996) for Mallorca, Tunisia, and Egypt; Mastronuzzi and Sansò (2000, 2004) for Apulia, Italy; Dominey-Howes et al. (2000) for the central Aegean; Minoura et al. (2000) for Crete and western Turkey, and Kelletat and Schellman (2002) for Cyprus. Italicized text indicates the principal tsunamigenic zones, including the Bay of Naples (BoN), Aeolian Islands (AI), Messina Straits (MS), Gulf of Corinth (GoC), Santorini (S), North Aegean trough (NAT), and Chios-Izmir zone (C-I). Mega-turbidites are from Rothwell et al. (2000) in the Balearic basin, and Hieke (2000) in the Ionian Sea. Deep-sea homogenite deposits in the eastern Mediterranean are from Cita et al. (1996) and Cita and Aloisi (2000). The lack of recorded tsunamis in some parts of the region, most notably the North African coast, undoubtably reflects a lack of records rather than a real absence of incidence. Data for the Atlantic and the Black Sea are not shown. See Chapter 17.

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TABLE 13.2. Amplitude, duration, permanence, and length of coast affected by various types of rapid relative sea-level change in the Aegean Process or type of change Local tsunami Storm surges Coseismic subsidence or uplift (moderate magnitude earthquakes) Semi-diurnal tides Marine flooding from barrier breaching (natural or artificial coastal barriers)

Amplitude Up to 30 ma 1 mb 0.3–0.8 m (this study) 1–3 m (Gulf of Corinth)c 0.6 m (Gulf of Evvia) Typically 0.1 m or lessd Decimetrese

Duration of change

Permanence of change

Length of coastline affected

Minutes Hours Seconds

15 dust storm days per year

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ranada Granad Granada Loes Loess

KO KONY NYA KONYA

TUNI TUNIS

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Main trajectories trajectorie of spring dus dust transport EL ARISH ARIS

Er TRIPOLI Grand Erg TRIPOL Orienta Oriental

Erg er Raoui Raou

CAIRO CAIR

Nege Negev Sina Sinai

Qattara Depressio Depression

Er Erg Iguidi

Er Erg

Er Erg ssaouane Issaouane

Edeye Calanscio Edeyen Calansci Ubar bari Ubari

0

500 km

Fig. 14.1. A map of some aeolian phenomena and locations in the Mediterranean basin.

Dust from North African sources has been collected from the atmosphere over the Mediterranean (Chester et al. 1984; Tomadin et al. 1984; Blanco et al. 2003) and several authors have noted its significant geological role as an input to Mediterranean deep-sea sediments (Tomadin 1974; Löye-Pilot et al. 1986; Tomadin and Lenaz 1989; Box et al. 2008). Guerzoni et al. (1999: 147) have suggested that ‘Both the magnitude and the mineralogical composition of atmospheric dust inputs indicate that aeolian deposition is an important (50 per cent) or even dominant (>80 per cent) contribution to sediments in the offshore waters of the entire Mediterranean basin.’ Aerosols may include appreciable iron concentrations that may influence marine biogeochemistry (Guieu et al. 2002). It has also been shown that dust inputs to the marine environment can stimulate bacterial activity in the water column. Pulido-Villena et al. (2008) have argued that the connections between dust inputs and the ocean carbon cycle may be more significant than previously thought. Dust may also be a primary source, through accumulation and weathering, for the formation of terra rossa soils on limestones (Yaalon and Ganor 1979; Macleod 1980; Rapp 1984),

though this is a controversial issue (Chapter 6). It is possible, for example, that limestone dissolution contributes to the clay fraction, though grain size and mineralogical analyses of insoluble residues are suggestive of the importance of aeolian inputs from distant or more local sources (van Andel 1998). The Saharan source strength for dust transport to Europe was estimated at 80–120 × 106 tons per year by D’Almeida (1986) based on sunphotometer readings taken in the early 1980s. A major source area for transport to western Europe was identified in southernmost Algeria, between the Hoggar Massif and Adrar des Iforhas (Figure 14.1) Another source, where material is particularly rich in palygorskite (Molinaroli 1996), is in western Sahara/southern Morocco. These sources have been essentially confirmed by back trajectory analysis for dust deposited over north-eastern Spain. Avila et al. (1997) traced deposition events back to three main areas: western Sahara, Moroccan Atlas, and central Algeria. An area south-east of Benghazi in Libya is also an important source region (Koren et al. 2003). Transport to southern Europe is frequent. A year of monitoring on Corsica, for example, revealed twenty

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dust events (Bergametti et al. 1989) originating in three source areas: eastern Algeria/Tunisia/western Libya, Morocco/western Algeria, and ‘south of 30◦ N . Dust transport has also been monitored through analysis of satellite borne Total Ozone Mapping Spectrometer (TOMS) data (e.g. Kalivitis et al. 2007). The basis for this method is described by Herman et al. (1997) and Hsü et al. (1999). Data for the sample year of 1999 show that dust penetrated the troposphere over the Mediterranean on more than 60 per cent of days throughout the months of March to September, with Mediterranean dust outbreaks recorded on 100 per cent of days in June and August (Table 14.1a) (Middleton and Goudie 2001). In the western Mediterranean, a swathe of Saharan material penetrated as far as Sardinia on 43 per cent of days in June and 60 per cent of days in August (Table 14.1b). The most frequent source for dust reaching Sardinia was central-southern Tunisia and adjacent areas of eastern Algeria, an area of salt pans or chotts shown by Dubief (1953) as generating more than forty vents de sable. Morocco/western Algeria was an TABLE 14.1. Dust over the Mediterranean Month

No. of dust events∗

Total observations

%

(a) Mediterranean basin January February March April May June July August September October November December Year

6 6 22 24 30 30 24 30 19 9 4 3

28 27 31 28 31 30 30 30 30 30 27 31

21.43 22.22 70.97 85.71 96.77 100.00 80.00 100.00 63.33 30.00 14.81 9.67

207

353

61.79

(b) Sardinia January February March April May June July August September October November December

1 0 3 4 11 13 9 18 4 2 0 0

28 27 31 28 31 30 30 30 30 30 22 31

3.57 0.00 9.68 14.29 35.48 43.33 30.00 60.00 13.33 6.67 0.00 0.00

Year

65

353

19.40

∗

Days with an Aerosol Index (AI) > 1.9.

Source: 1999 TOMS data in Middleton and Goudie (2001).

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occasional source of dust reaching Sardinia according to the 1999 TOMS data, as was north-eastern Libya, the latter a source not mentioned by Bergametti et al. (1989) in their Corsican work. Analysis of meteorological station data by Middleton (1986) showed a broad area across eastern Algeria and western Libya with >15 dust storm days a year on average and three stations in north-eastern Libya with a long-term (1956–77) mean of >10 dust storm days a year. According to TOMS, the least active months in 1999 for the whole Mediterranean were November and December, both showing less than 15 per cent of days with material over some part of the Mediterranean. No outbreaks reached Sardinia in November, December, or February, and occurred on just one day in January (Middleton and Goudie 2001). Since the source strength and transport of Saharan dust to Europe is heavily dependent upon climatic parameters, variations in dust fall frequencies could be an indicator of climatic change and this aspect has attracted the attention of several studies in recent years. However, year-to-year variability is high, as shown by eleven years of deposition records (1984–94) for Corsica where the annual input of Saharan dust varied between 4.0 and 26.2 g m−2 over this period (Löye-Pilot and Martin 1996). Nonetheless these authors noted that annual deposition rates over Corsica were seen to peak in the 1980s, and declined in the first half of the following decade. Deposition over Corsica shows a bimodal seasonality (spring and autumn maxima) similar to that recorded in north-eastern Spain where a decline in the frequency of red rain events in the 1990s has been noted after a peak of activity in the late 1980s (Avila et al. 1997; Avila and Peñuelas 1999). Several other studies have remarked upon the peak in Saharan dust falls over Europe in the late 1980s. Dessens and Van Dinh (1990) noted a marked increase in the frequency of Saharan dust outbreaks depositing at the Midi-Pyrenees Aerology Observatory in Lannemezan, France, over the period 1983–9, a rise confirmed from a much longer dataset (1949–94) from the Spanish Mediterranean coast south of Alicante (Sala et al. 1996). The long-term average there was approximately 2 dust-rain days per year, but from 1985 to 1994 the annual total averaged 6.5 dust-rain days, with 9.0 dust-rain days per year recorded for the period 1989–94. A significant increase in the quantities of Saharan dust falling over the French Alps since the early 1970s, with very high inputs occurring after 1980 (De Angelis and Gaudichet 1991), was detected from an ice core that yielded dust deposition data for a

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thirty-year period (1955–85). Contrary to this evidence, however, Conte et al. (1996) show a decline in the frequency of strong siroccos over the period 1951–90 at Trapani in Sicily, probably due to an increase in anticyclonic activity in the western and central parts of the Mediterranean basin which tends to counteract the occurrence of frontal disturbances that generate the strong, dust-laden southerly winds from the Sahara (Chapter 3).

Eastern Mediterranean Trajectories Transport from North Africa to the eastern Mediterranean occurs predominantly during spring and is commonly associated with the eastward passage of frontal low pressure systems, while dust from sources in the Middle East is more typically transported to the Mediterranean in the autumn (Dayan 1986; Kubilay et al. 2000). Analysis of 23 heavy dust falls in Israel over a 20-year period suggest that the North African type is by far the most common (Ganor et al. 1991), and North African dust is also the main source in Turkey (Mace et al. 2003). Long-range transport of dust to the eastern Mediterranean from the Arabian Desert also occurs in events that tend to last for one day, whereas the transport of Saharan dust to the central Mediterranean is characterized by events lasting 2–4 days (Dayan et al. 1991). There is some seasonal variation in the source areas of dust reaching Israel, with Chad being the spring source, Egypt and the Red Sea the source in July/August, and Libya in the autumn (Israelevich et al. 2003). Central Algeria is identified by Ganor et al. (1991) as the most frequent source area for Saharan dust reaching Israel, while Ganor and Foner (1996) distinguish between material commonly transported from sources in the Hoggar and Tibesti mountains in northern Chad, the latter also picking up material from the Western and Sinai Deserts. Some 25 × 106 tons of Saharan dust were estimated by Yaalon and Ganor (1979) to reach the eastern Mediterranean basin annually, most settling into the Mediterranean Sea. This figure has subsequently been revised upward, to 70 × 106 tons per year (Ganor and Mamane 1982) and more recently to 100 × 106 tons per year (Ganor and Foner 1996). The increase reflects the steady rise in frequency of Saharan dust episodes over Tel Aviv from ten per year in 1958 to nineteen per year in 1991 (Ganor 1994). The passage of Saharan dust into the Levant is of relatively frequent occurrence. The type of synoptic situation that leads to this is the passage of an advancing cold

front and the associated strong surface winds ahead of it penetrating south-eastwards from the Mediterranean Sea deep into the northern Sahara and Libyan Desert (Michaelides et al. 1999). This can be illustrated by the situation in mid-March 1998 (Goudie and Middleton 2002), when dust proved to be a significant natural hazard. A major dust event caused ports and airports to be closed, created breathing problems for inhabitants of Amman in Jordan, and led to fatal motoring accidents in Egypt and Jordan. Mean visibility in Amman was reduced to 4.2 km. A large, deep depression moved eastwards from North Africa, and then deepened further over the Middle East as it encountered cold polar air pushing across Turkey. The progress of this system can be traced by looking at the daily TOMS maps for the period from 14–19 March (Figure 14). The sequence starts with an area in eastern Algeria, southern Tunisia, and north-western Libya generating Aerosol Index (AI) values greater than 2.6. This index is a semi-quantitative measure of the amount of dust aerosol in the atmosphere, with higher index values generally indicating higher dust loadings. Clean air has a value of 0 and as a rule of thumb dusty air has values that exceed 2.0. The following day the dust pall has moved across Libya into Egypt and the eastern Mediterranean. On 16 March the main area with high AI values covers Cyprus, Israel, Syria, Jordan, and Sinai. Mean visibility at Amman airport was reduced to 3.2 km. On 17 March the area with high AI values has broken down into a group of small, deep clusters and by 19 March most of the area has values less than 1.0. Yet another example of the movement of a dust storm from the northern Sahara to the Middle East is provided by the events of April 2000 (Figure 14.3). On 18 April TOMS shows a large area of dust over southern Libya and far western Egypt. On the following day this reached eastern Egypt, Israel, and Lebanon, causing the closure of the port of Alexandria and the cessation of flights to Aswan. The mean visibility at Cairo airport was reduced to 4.6 km and that at Luxor to 3.7 km. In Israel mean visibility in Beersheva was reduced to 2.6 km. Limassol in Cyprus was also badly affected, as were flights in southern Turkey. The dust event than moved over Jordan, where the level of activity was less. Rates of dust deposition decline with distance from source (Table 14.2). Thus the annual values for western Europe (e.g. Central France and the Alps) are less Fig. 14.2. (opposite) The passage of dust systems from North Africa to the Middle East, mid March 1998, based on TOMS (from Goudie and Middleton 2002). The scale shows Toms Aerosol Index (AI) values.

14/3/98

15/3/98

16/3/98

17/3/98

18/3/98

19/3/98

1.0 1.4 1.8 2.2 2.6 3.0 3.4

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Andrew Goudie 17/4/00

18/4/00

19/4/00

0.7

1.1

1.5

1.9

2.3

2.7

3.1

TOMS Aerosol Index (AI) High ground

Fig. 14.3. The TOMS sequence across North Africa to the Middle East for mid April, 2000 (from Goudie and Middleton 2002).

then 1 g m−2 . Further south, in north-east Spain, a value of 5.1 g m−2 is recorded, while over Sardinia, Corsica, Crete, and the south-east Mediterranean, most values are between 10 and 40 g m−2 . Included in the dust deposits may be such clay minerals as palygorskite (Molinaroli 1996; Tomadin et al. 1984) and kaolinite (Foucault and Mélières 2000; Avila et al. 1996). These may reflect different source areas, with the former tending to imply a more arid origin than the latter. Krom et al. (1999) used strontium isotopes as a tracer to investigate the relative contribution of Nile suspended sediments and Saharan dust inputs to the eastern

Mediterranean Sea, while Frumkin and Stein (2004) used strontium and uranium isotopes to identify Saharan dust in speleothems in Jerusalem. Grain size data for dust are provided in Table 14.3. These indicate that most of the dust in the Mediterranean is fine to medium silt.

Loess As already mentioned, dust deposition has contributed to the nature of both terrestrial and marine sediments. It has also been found in lake sediments (Narcisi 2000). The role of dust is especially important in the case

Aeolian Processes TABLE 14.2. Dust deposition amounts across the Mediterranean Source

Nihlen and Olsson (1995) Le-Bolloch et al. (1996) Wagenbach and Geis (1989) De Angelis and Gaudichet (1991) Avila et al. (1996) Bergametti et al. (1989) Löye-Pilot et al. (1986) Bücher and Lucas (1984) Pye (1992) Herut and Krom (1996) Herut and Krom (1996) Hernández and Hernández (1997)

Location

Aegean Sea Southern Sardinia Swiss Alps French Alps NE Spain Corsica Corsica Central France Crete Israel coast SE Mediterranean SE Spain

Annual deposition (g m−2 ) 11.2–36.5 6–13 0.4 0.2 5.1 12 12.5 1 10–100 72 36 23

of aeolian silts (loess), deposits of which are known from various parts of the Mediterranean basin (CoudéGaussen 1990). The most notable deposits are probably those from southern Tunisia (Figure 14.4) (Brunnacker 1973;

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TABLE 14.3. Particle size characteristics of dust in various parts of the Mediterranean Source

Coudé-Gaussen (1991) Mattsson and Nihlen (1996) Sala et al. (1996) Pye (1992) Ozer et al. (1998) Bücher and Lucas (1994) Coudé-Gaussen (1991) Tomadin et al. (1984) Blanco et al. (2003)

Location

Modal, mean or median size in microns (Ïm)

Maghreb Crete Spain Crete Genoa, Italy SW France South of France Central Mediterranean Southern Italy

5–40 (median) 8–30 (modal) 4–30 (mean) 4–16 (median) 14.6 (median) 4–2.7 (median) 8–11 (median) 2–8 (modal) 1.7–2.4 (median)

Coudé-Gaussen et al. 1982), where the Matmata plateau loess reaches a thickness of 18 m at Téchine and contains up to five palaeosols typically rich in smectite and palygorskite. The loess probably derives from the Chott Djerid (a large ephemeral lake to the west) and from the Grand Erg Oriental (a large sand sea immediately to the

Fig. 14.4. The loess of Matmata, southern Tunisia, has been inhabited by cave dwellers (photo: Andrew Goudie).

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south-west). Coudé-Gaussen et al. (1983) suggest that two great phases of deposition occurred between 28,000 and 10,000 BP and from 6,000 to 4,000 BP, and CoudéGaussen et al. (1983) believed that maximum loess deposition occurred during humid conditions. This is a view that was disputed by Dearing et al. (1996) following their investigations of the mineral magnetic properties of these deposits. They believed that the period between 15,000 and 20,000 years BP was a time both of aridity and accelerated loess deposition. More recently, using luminescence dating, Dearing et al. (2001) have shown that some of the loess is older than this, with a sequence of loess and palaeosols from Téchine being deposited during the period 100,000–250,000 years BP. Loess has been known from Libya for many years (Schwegler 1944). Two loess deposits in Libya have also been recently studied. These have been classified as silty loess in the Tripoli region in the north-west of the country and clayey loess in the Ghat area in the southwest (Assallay 1996). Other significant loess deposits occur in Sinai (Rögner and Smykatz-Kloss 1991), but the most extensive deposits in the Middle East are those of the Negev (Yaalon and Dan 1974). Late Pleistocene loess, up to 10 m thick, occurs in the Granada basin of south-east Spain (Günster et al. 2001). Other loess is known from the central Apennines of Italy (Frezotti and Giraudi 1990), the Po Valley (Busacca and

Cremaschi 1998; Castiglioni 2001), Susak Island in the Dalmatian Archipelago (Cremaschi 1990), and in parts of Greece, including Crete (Brunnacker 1980). In most cases it is still not clear just to what degree the different loess deposits are derived from local or more distant processes, and to what extent some are derived from deserts and ephemeral lakes and some from formerly glaciated areas with large braided stream systems. Recent work in the Negev of southern Israel by Crouvi et al. (2008) has argued that the coarse silts of hilltop loess sequences were derived from the sand seas that advanced into Sinai and the Negev during the Late Pleistocene when the Mediterranean shelf was exposed.

Dunes and Other Aeolian Forms Because of the dryness of parts of the Mediterranean region, sand dunes have formed in various parts of the region, including its coastlines and arid interiors. Aeolian sand has also accumulated in various caves, providing important information about palaeoenvironments associated with archaeological sites (see e.g. Farrand 1979), especially in coastal settings. Some of the sand deposits may be the product of past more arid conditions, as is the case with the Plio-Pleistocene sands

Fig. 14.5. Barchan dunes in the Libyan Desert, Kharga depression, Egypt (photo: Andrew Goudie).

Aeolian Processes

of the Valdarno Basin of the northern Apennines in Italy (Ghinassi et al. 2004). Likewise, in Mallorca, dune sand appears to have been transported from exposed carbonate shelves far inland by westerly winds under cold, dry, windy phases of the Pleistocene (Nielsen et al. 2004). One extensive dune area occurs on the Mediterranean coast of Sinai, in the vicinity of El Arish. The so-called Sinai–Negev erg, which covers about 10,000 km2 , is formed of sand derived from the Nile, which has been reworked by the Mediterranean and blown inland from its beaches (Tsoar and Goodfriend 1994), and from the Mediterranean shelf at times of lowered sea level (Crouvi et al. 2008). Much of this erg, which contains both transverse and linear forms, developed between 19,000 and 10,000 years BP (Goring-Morris and Goldberg 1990), but there have been various phases of remobilization in the Holocene that could have resulted from anthropogenic pressures. Indeed, many of the coastal parabolic dunes of the Israeli and Palestine coast have formed

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in the last 1,000 years, probably because of vegetation removal associated with agricultural and pastoral activities (Tsoar and Blumberg 2002). Only 17 per cent of the Israeli coastal dunes are still of good or reasonable ecological value (Kutiel 2001), and recreational impacts can be severe (Kutiel et al. 1999). Elsewhere in the Near East, some Late Pleistocene and Holocene dunes have been studied in the Konya plain of Turkey (Kuzucuoglu et al. 1998). Inland dunes are also known from the Po valley in northern Italy (Castiglioni 2001). Coastal dune systems are extensively developed around the Mediterranean, not least in the Balearic Islands, where vegetated parabolic forms are widespread (Servera and Rodriguez 1996), together with cliff-front ramps (Clemmensen et al. 2001). However, large tracts of them are being modified or destroyed by tourist pressures (Schmitt 1994; Chapter 13). In mainland Spain parabolic dunes are also common in Catalonia (Gos and Serra 1993) and Valencia (Sanjaume

Fig. 14.6. Gypsum crust soils with polygonal structures in southern Tunisia. Deflation of gypsum from dry lake basins and its deposition downwind contributes to the formation of these widespread surface material types (photo: Andrew Goudie).

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and Pardo 1991, 1992), but are also being subjected to intense human pressures. Perhaps the most celebrated dunes of Spain are, however, on its Atlantic coast in the Doñana National Park (Munoz-Reinoso 2001), but here too human activities (including over-exploitation of groundwater) pose a threat. A general review of sanddune management problems and techniques in Spain is provided by Gómez-Pina et al. (2002). France also has many dune systems along the coast of the Gulf of Lions (Corre 1991). Urban sprawl is a major threat to coastal dunes along the coast of Egypt, especially between Alexandria and El Alamein (Batanouny 1999). One of the most interesting features of many of the coastal dunes on the shorelines of the Mediterranean is the formation of lithified carbonate-rich dunes or aeolianites. These occur in mainland Spain, the Balearic Islands, Corsica, and Sardinia, Greece (Gittenberger and Goodfriend 1993), Cyprus and Crete (Kelletat 1991), North Africa (El-Asmar 1994; Plaziat and Mahmoudi 1990), and along the Levant coast. A bibliography is provided by Brooke (2001). In Israel this material is often called kurkar and contains multiple

palaeosols (hamra) indicating alternating stability and dune activity (Verrecchia 1989; Tsoar 2000; Frechen et al. 2002) during the Late Pleistocene (Sivan and Porat 2004). On the south side of the Mediterranean extensive dunes occur in Morocco, Algeria, Tunisia, Libya, and Egypt—the northern Saharan dunes. They form great sand seas, which from west to east are the Grand Erg Occidental, the Grand Erg Oriental, the Calanscio Erg, the Great Sand Sea, and the Abu Moharik belt (Goudie 2002) (Figures 14.1 and 14.5). These northern Saharan sand seas have complex patterns of dunes as a consequence of multidirectional wind regimes involving the interaction of both the northeasterly trades and the mid-latitude circulation (Chapter 3). It is, of course, difficult to say where the Mediterranean region stops and the Sahara starts, but undoubtedly large tracts of the northern Sahara have important impacts on the Mediterranean lands. In Tunisia and Algeria the presence of a series of closed depressions—the Chotts—which may themselves in part be the product of aeolian deflation from

Fig. 14.7. The great lunette dune on the lee side of the Sebkha el Kelbia, central Tunisia. The salt lake from which it is derived is visible in the background (photo: Andrew Goudie).

Aeolian Processes

structurally controlled basins (Sweezey 1996, 1997)— has promoted the development of gypsum dunes and has contributed to the formation of some gypsum crusts (Figure 14.6) (Watson 1985) and of some enormous lunettes on their lee sides (Coque 1979; Perthuisot and Jauzein 1975; Goudie and Wells 1995; Benazzouz 1986). In the case of the Sebkha el Kelbia in Tunisia the lunette is 146 m high (Figure 14.7). In Egypt, just to the south of the Mediterranean coast, is the Qattara Depression, the base of which lies at −133 m below present sea level. The origin of this 19,500 km2 depression has been the subject of considerable controversy. Some workers have seen it as essentially deflational, others as being tectonic, others as being related to old river courses, and others as being the result of a combination of deflation and salt weathering (Goudie 2002; Aref et al. 2002; Ball 1927). Another area of closed drainage depressions in the Mediterranean region is the Ebro basin of Spain (mean annual rainfall 350 mm). Here there are a large number of closed depressions, called ‘saladas’. Their origin partly results from solution of Tertiary evaporite strata, but they have been deepened and extended by deflation (Sánchez et al. 1998). Yardangs (wind fluted rock) and nebkha dunes (which accumulate around vegetation) have developed in the same area (Gutiérrez-Elorza et al. 2002). Interestingly, clay dunes, derived from deflated saline playas, occur in the Guadiana valley of Spain (Rendell et al. 1994). Thus we have the whole spectrum of aeolian particle sizes in dune systems in the Mediterranean region. Finally, wind action plays a major role in the dispersal of volcanic tephra across the region (Chapter 15) and its subsequent deposition on land and in the sea (Narcisi and Vezzoli 1999) (Chapter 2). It is also responsible for the transport of large quantities of anthropogenic atmospheric pollutants into the area from Europe, Asia, and beyond (Lelieveld et al. 2002) (Chapter 22).

Conclusion Saharan dust has a major impact on the climate, biogeochemical cycling, atmospheric quality, and soils of the Mediterranean countries. It also constitutes a significant natural hazard. In the Pleistocene, the Sahara and other sources led to the development of scattered but locally extensive loess deposits, most notably in Tunisia and the Negev. Also important has been the development of coastal dunes, many of them now under severe anthropogenic pressures, and of larger continental dune

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systems, especially in North Africa and the Negev/Sinai region. Wind erosion has created erosional features such as deflation basins and yardangs in North Africa and semi-arid Spain. Wind erosion and dune degradation are being exacerbated at the present time by burgeoning tourism and recreation, by deforestation, cultivation, overgrazing, and the lowering of water tables, but overall it is not as potent a threat as that posed by water erosion and slope destabilization.

References Aref, M. A. M., El-Khoriby, E., and Hamdan, M. A. (2002), The role of salt weathering in the origin of the Qattara Depression western Desert, Egypt. Geomorphology 45: 181–95. Assallay, A. M., Rogers, C. D. F., and Smalley, I. J. (1996), Engineering properties of loess in Libya. Journal of Arid Environments 32: 373–86. Avila, A. and Peñuelas, J. (1999), Increasing frequency of Saharan rains over northeastern Spain and its ecological consequences. The Science of the Total Environment 228: 153–6. Queralt, I., Gallart, F., and Martin-Vide, J. (1996), African dust over northeastern Spain: mineralogy and source regions, in S. Guerzoni and R. Chester (eds.), The Impact of Desert Dust across the Mediterranean. Kluwer, Dordrecht, 201–15. Queralt-Mitjans, I., and Alarcón, M. (1997), Mineralogical composition of African dust delivered by red rains over northeastern Spain. Journal of Geophysical Research 102 (D18): 21,977–96. Ball, J. (1927), Problems of the Libyan Desert. Geographical Journal 70: 21–8, 105–28, 209–24. Batanouny, K. H. (1999), The Mediterranean coastal dunes in Egypt: an endangered landscape. Estuarine Coastal and Shelf Science 49 (Suppl. A): 3–9. Benazzouz, M. (1986), Recherches géomorphologiques dans les hautes plaines de l’est Algérien: la sebkhat Tarf (Algérie). Ph.D. Thesis. University of Paris I, Sorbonne. Bergametti, G., Gomes, L., Remoudaki, E., Desbois, M., Martin, D., and Buat Ménard, P. (1989), Present transport and deposition patterns of African dusts to the north-western Mediterranean, in M. Leinen and M. Sarnthein (eds.), Palaeoclimatology and Palaeometeorology: Modern and Past Patterns of Global Atmospheric Transport, NATO ASI Series C. 282: 227–52. Blanco, A., De Tomasi, F., Filippo, E., Manno, D., Perrone, M. R., Serra, A., Tafuro, A. M., and Tepore, A. (2003), Characterization of African dust over southern Italy. Atmospheric Chemistry and Physics 3: 2147–59. Box, M. R., Krom, M. D., Cliff, R., Almogi-Labin, A., Bar-Matthews, M., Ayalon, A., Schillman, B., and Paterne, M. (2008), Changes in the flux of Saharan dust to the East Mediterranean Sea since the last glacial maximum as observed through Sr-isotope geochemistry. Mineralogical Magazine 72: 307–11. Brooke, B. (2001), The distribution of carbonate eolianite. EarthScience Reviews 55: 135–64. Brunnacker, K. (1973), Einiges über Löss-Vorkommen in Tunesien. Eiszeitalter und Gegenwart 23/4: 1–11.

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This chapter should be cited as follows Goudie, A. S. (2009), Aeolian processes and landforms, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 415–429.

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III

Hazards

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Editorial Introduction Jamie Woodward

Catastrophic earthquakes, explosive volcanic eruptions, and devastating storms and floods are intimately bound up within the history and mythology of the Mediterranean world. It is a key region for the study of natural hazards because it offers unrivalled access to long records of hazard occurrence and impact through documentary, archaeological and geological archives. Early texts and archaeological data have provided unique insights into the nature and impact of past eruptions, earthquakes, tsunamis, and other hazards. Notable events were carefully documented in Antiquity and the archaeological record provides insights into the impact of catastrophic events on past human societies. The eruption of Vesuvius in AD 79, for example, was famously documented by Pliny the Younger, and the excavations at Pompeii have provided extraordinarily rich insights into the dynamics and impacts of tephra falls and pyroclastic flows. The significance of environmental hazards in the demise of civilizations such as Minoan Crete (tsunami) and the Early Bronze Age in the Near East (drought) has been vigorously debated for decades. While such events have undoubtedly threatened people in the region since prehistoric times, the actual threat to human society has increased dramatically in the historical and modern periods as urban environments and their populations have rapidly expanded. This part of the volume analyses hazards associated with both endogenic and exogenic Earth processes and the interactions between them. It includes volcanic processes, crustal instability, tsunamis, fluvial floods, extreme weather phenomena, and wildfires. Each chapter explores the basic controls and the geography of a particular hazard and related processes, and, over a range of timescales, magnitude and frequency relationships and the nature of the threat to human society.

High-magnitude events are a fundamental part of the physical geography of the Mediterranean and play a key role in long-term landscape evolution and ecosystem change. Even though the processes associated with each hazard typically take place over very short timescales, they can set in motion longterm adjustments to geomorphological and ecosystem processes. Tephra falls can change soil properties and vegetation communities, for example, and earthquakes may trigger base-level change and landslips in river basins that enhance fluvial sediment yields for many centuries. The Mediterranean has a distinctive meteorology that can produce extreme weather phenomena. A range of weather-related hazards has led to major losses of life on land and at sea. River basins in the region are characteristically small, with steep channels that have always been prone to high discharges. Yet even though flood magnitude decreased in some areas during the twentieth century, the flood hazard increased markedly as urban areas expanded across valley floors and along coastal plains. Wildfires devastated large areas of Greece in 2007 and climate change predictions for the twenty-first century suggest they will become an increasing problem across the Mediterranean. Other significant hazards are considered elsewhere in this volume and include storm surges and coastal flooding, mass movements, collapse features in karst terrains, dust storms, prolonged droughts, and heat waves. Each chapter in Part III examines hazard mitigation strategies and explores the problems that each hazard poses for Mediterranean society. Recent decades have seen the wider application of remote sensing and GIS tools for hazard zoning and increasingly effective collaboration between Mediterranean countries in systems of monitoring and warning.

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15

Volcanoes Clive Oppenheimer and David Pyle

Introduction The historical record of Mediterranean volcanism is arguably the richest available for any region of the world. Documentary records date back to the Classical period, and archaeological records date back further still (Stothers and Rampino 1983; Chester et al. 2000). The Mediterranean is also home to some of the most famous, or indeed infamous, volcanoes on Earth (Table 15.1), several of which still present major threats to society today (Kilburn and McGuire 2001; Chester et al. 2002; Guest et al. 2003). A number, for example Santorini, Etna, and Vesuvius, have menaced human populations since Antiquity, and the human response and risk perception today are strongly shaped by a culture, which itself owes much to the volcanic landscapes and eruptions (Chester et al. 2008). The science of volcanology was born, and has since flourished, in the cradle of the Mediterranean. It began, arguably, with the careful descriptions by Pliny the Younger of the AD 79 eruption of Vesuvius that buried Pompeii, and developed through the scientific investigations of Sir William Hamilton in the eighteenth century. The region was the playground of the pioneers of modern volcanological studies in the nineteenth century (e.g. Fouqué 1879), and today it boasts a number of state of the art volcano observatories such as that which monitors Vesuvius. Several volcanoes, eruption styles, geothermal manifestations, and rock types have inspired nomenclature now widely used within the volcanological community: plinian, vulcanian, and strombolian eruptions; low temperature gas emanations known as solfataras; rocks known as pantellerites. The very word ‘volcano’ comes from the Aeolian island Vulcano, where Vulcan’s forge was situated. The sea-filled crater of Santorini was one of the

first volcanic ‘calderas’ to be described, by Ferdinand Fouqué in the 1870s.

General Geological Setting for Volcanic Activity The Mediterranean basin tracks the geological suture between the African plate to the south, and the Eurasian tectonic plate to the north (Chapters 1, 13, and 16). Many regions along this suture have experienced volcanic activity within the past 10–20 Myr (million years), most of it related to the continuing process of subduction that has consumed the northern margin of the African plate. Geologically, this has been a very messy collision, and our understanding of why particular volcanoes are found in particular places remains incomplete. During the past million years or so, the regions of active volcanism around the Mediterranean have become settled into the pattern that we see today (Figure 15.1). The active volcanic regions are more or less confined to central and southern Italy (from the Campanian and Roman provinces of the mainland, to the Aeolian Islands, to Etna, on Sicily, and Pantelleria; Figure 15.1); and to an arc of volcanism in the southern Aegean sea (from Aegina and Milos in the west, to Kos, Nisyros, and Yali in the east; Figure 15.1). Inland and further to the east, the trail of volcanism stretches through the volcanic province of Cappadocia, in Turkey, past the fabled, but poorly known, Ararat, and up into the great volcanoes of the Caucasus: Elbrus and Kazbek. In addition to the volcanoes that flank parts of the Mediterranean, mainland Europe has also seen geologically young episodes of volcanic activity, most recently as the ice sheets retreated, in regions including the

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Clive Oppenheimer and David Pyle TABLE 15.1. Major ash layers correlated with known volcanic eruptions in the Mediterranean over the past 200 kyr

Eruption age (kyr BP)1 3.6 16.5 22.0 39.3 45 150 161 180 203

Eruption name

Other identifier (e.g. ash layer)

Source volcano

Composition

Magma volume (km3 )

Minoan Biancavilla-Montalto Cape Riva Campanian Ignimbrite Green Tuff Middle Pumice Kos Plateau Tuff Lower Pumice 2 Lower Pumice 1

Z2 Y1, Et1 Y2 Y5 Y6 W2 W3 V1 V3

Santorini Etna Santorini Phlegrean fields, Italy Pantelleria Santorini Kos Santorini Santorini

Rhyodacite Sodic andesite Rhyodacite Trachyte Peralkaline rhyolite Dacite Rhyolite Rhyodacite Rhyodacite

28–31

105–210 (>10) (>60) (>10) (>10)

Minimum tephra coverage area (km2 )

Additional references2

3 × 105 4 × 104 3 × 105 2–4 ×106 7 × 104

a b, c d e f

104 105 1.5 × 105

g

1

Eruption ages are ‘best estimates’ based on isotopic age constraints (calibrated 14 C or 40 Ar/39 Ar) or correlation with orbitally tuned marine core timescales. The correlation of V3 is newly suggested here. Additional references: (a) Pyle (1990); (b) Paterne et al. (1988); (c) Coltelli et al. (2000); (d) Wulf et al. (2002); (e) Pyle et al. (2006); (f) Civetta et al. 1988; (g) P. E. Smith et al. (1996). 2

Source: Compiled from data in Keller et al. 1978; Narcisi and Vezzoli 1999; Druitt et al. 1999; Lourens 2004.

Europe, which decompressed and melted during postglacial ice removal and flexuring (Nowell et al. 2006).

Auvergne and Massif Central of France; the Eifel region and the Rhine valley (Germany); around the flanks of the Pannonian basin of Hungary, Slovakia, and the Czech republic; and the Olot region of northern Spain (Scarth and Tanguy 2001). Geological explanations for the origin of these magmas are not agreed, but may relate to patterns of active rifting or to the presence of anomalously ‘warm’ regions of the mantle beneath

Volcanic Hazards Volcanoes can pose hazards between, as well as during, eruptions (Table 15.2). Volcanic disasters have often occurred months or years after eruptions begin, striking

Eifel Eifel

Chaine des Chaine desPuys Puys

Caucasus

Tuscany

Elbrus

Kazbek

Roman Province

Olot Olot

Campanian Province

Ararat

Amiata Colli Albani Ischia

Aeolian

Cappadocia Vesuvio

Stromboli i Lipari Vulcano Etna

Pantelleria Pntelleriaa

Campi Flegrei CampiFlegrei Mar di Sicilia

Kula Methana Methana Milos Santorini

Kos Kos Yali Nisyros

Aegean Volcanoes which have erupted since 1900 Other named volcanoes Other volcanoes

0

500 km

Fig. 15.1. Map of the Mediterranean basin showing the locations of selected volcanoes and volcanic provinces discussed in the text.

Volcanoes

437

TABLE 15.2. Summary of volcanic hazards Hazard type Pyroclastic density currents Debris avalanches, debris flows, lahars Tsunamis Airborne ash

Tephra falls

Landslides (debris avalanches) Lava flows

Point-source and diffuse gas and acid particle emissions (notably SO2 , sulphuric acid aerosol, HCl, HF, CO2 , H2 S, and radon)

Acid rain (due to rainfall through volcanic clouds) Ballistics (bombs, blocks) Earthquakes Ground deformation (e.g. subsidence, ground-cracking) Atmospheric shock waves Lightning in volcanic clouds Volcano restlessness or threat of eruption

Elements at risk, consequences Highly destructive to property and buildings, pyrogenic, cause thermal injury, and associated with a very low survival rate amongst exposed individuals. Drowning, impact injuries, extensive property destruction. Human health, property. Respiratory and cardiovascular health hazard (asthma, bronchitis, pneumoconiosis) and an irritant to eyes and skin, affects aircraft operations; impairs, disrupts, or prevents telecommunications. Damage to property, transport systems, human and animal health, water quality (corrosion of buildings; contamination of water supplies by volatiles, e.g. fluorine, carried on ash or blocking water treatment filters; impacts of agricultural/farming losses due to fluorine, burial under ash, etc.) Damage to health, property. Highly damaging to property, roads, agricultural land, forests, etc.; generally sluggish enough for people to evade except in some particular cases where lavas are very fluid and fast moving, or start uncontrolled fires. Impacts on air quality and human health, vegetation health and soil condition (agriculture). Acid gases cause bronchoconstriction, aggravation of respiratory disease; CO2 & H2 S: asphyxiation, possible neurological problems, nasal and lung cancers; radon creates lung cancer risk with long-term exposure. Chemical conjunctivitis and respiratory effects associated with HCl-rich clouds (known as LAZE for lava-haze) formed when lava enters the sea. Vegetation health (agriculture), property, water quality; irritant to eyes, skin. Impact injuries, burns. Property damage. Property damage and resulting physical injuries. Structural damage to property, roads. Property damage. Danger of electrocution. Economic impacts to regional development, tourism, investment, etc.

Source: Modified from Francis and Oppenheimer (2004) and Hansell and Oppenheimer (2004).

when the population at risk has become inured to the threat (Simkin et al. 2001). The intensity and magnitude of eruptions need not scale with their impacts, since exposure and vulnerability of elements at risk vary tremendously from one volcano to another. While the history of a volcano’s activity (witnessed in the rock record) points to the nature of potential future behaviour, the eruption style—and consequent hazards— are highly variable from one eruption to the next, and even within a single eruptive episode. Between eruptions, volcanoes may emit harmful gases and particles, and continue to pose serious hazards in the guise of mudflows, debris avalanches, and volcanogenic tsunamis. In a recent review of the literature, Witham (2005) found that 490 volcanic events in the twentieth century resulted in human impacts, with 4–6 million people evacuated, made homeless, or otherwise affected. Fatalities occurred in around half the events, with an estimated total death toll of 80,000–100,000. The risk of catastrophic human and economic losses from future eruptions is significant, especially given the spread of

urbanization in many volcanic regions—perhaps most notably in the Campanian district of Italy that includes the city of Naples. A further notable feature of the statistical record is that the number of injured (about 12,000) is much lower than the number of deaths (>80,000)— volcanic phenomena are often associated with low survival rates in impacted areas. Post-eruption famine and epidemics, pyroclastic density currents, mudflows, and volcanogenic tsunamis account for the majority of recorded deaths arising from volcanism. Few deaths have been associated with lava flows, however. While they may result in considerable damage to infrastructure and property they usually move sufficiently slowly to permit the evacuation of residents. The Mediterranean has escaped any major explosive activity since the early 1900s, when eruptions of Vesuvius in 1906 caused major economic impacts in Naples. One effect of these events was relocation of the 1908 Olympic Games, scheduled to take place in Rome, to London. Funds that were to have been used to support the Games in Rome were diverted to the reconstruction efforts in Naples. The following sections review a range

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Clive Oppenheimer and David Pyle

(a)

(b)

Fig. 15.2. Volcanic hazards. (a) Oblique aerial view of Mt. Etna erupting in 2001. The dense ash cloud rises from the Laghetto cinder cone, formed during this eruption. (b) The 2001 eruption of Mt. Etna emplaced lava flows that destroyed part of the tourist village on the southern approach to the mountain. Fumes rise from the active margin of the flow. The tourist village has been protected in part by use of earth barriers but the road has been crossed by the new lavas (photos: Clive Oppenheimer).

of volcanic processes, products, and related hazards and are followed by a detailed examination of these phenomena from a Mediterranean perspective.

Tephra Clouds and Falls Ash clouds are associated with explosive eruptions (Figure 15.2a) and can pose a risk to aviation through damage to jet engines and abrasion of cockpit windows. To date there have been no reported air crashes arising from encounters with volcanic clouds, but there have been several near misses that resulted in substantial damage to the aircraft involved. Tephra fallout presents several kinds of hazard: by causing buildings to collapse due to loading on roofs and by impacts on terrestrial and aquatic environments. Heavy falls of tephra can damage vegetation, including agricultural crops, and pose a respiratory health hazard to humans. Tephra can also carry significant quantities of volatiles, including fluorine, scavenged from the eruption plume. Grazing animals can ingest toxic quantities of fluorine when ash contaminated in this way lies on the ground (Cronin et al. 2003). It may also lead to corruption of drinking water with the potential to cause fluorosis in human populations.

Pyroclastic Density Currents (PDCs) One of the most life-threatening aspects of volcanic activity results from exposure to clouds of hot gas and tephra known as pyroclastic density currents (PDCs; Branney and Kokelaar 2002). Past PDC events have resulted

in dead:injured ratios of 10:1 or greater—significantly higher than in any other type of natural disaster. PDCs typically form during the collapse of an explosive eruption column (e.g. Vesuvius in AD 79) or from gravitational failure of a lava dome (as witnessed on countless occasions at Soufrière Hills volcano, Montserrat, over the past decade). They can be envisaged as hot hurricanes of ash, rock, and gases, and can travel at speeds of 350 km hr−1 , or more, and reach temperatures of several hundred degrees Celsius. During an eruption of Mont Pelée on Martinique in 1902, PDCs swept into the town of St Pierre at an estimated speed of 160 km hr−1 , resulting in the deaths of 29,000 people within minutes (including all but two of the city’s population). The main causes of death are heat-induced shock, asphyxia, thermal lung injury, and burns. Survivors from PDCs tend to have been exposed to only the more dilute parts of the current or sheltered in some way (Loughlin et al. 2002), but can be critically injured due to respiratory and skin burns. Prior evacuation from areas at risk from PDCs is the only recommended way to minimize fatalities, for example through the establishment of exclusion zones. Many of the fatalities in Pompei and Herculaneum in AD 79 were believed to be from PDCs.

Lava Flows The term ‘lava flow’ usually refers to erupting lava that has the opportunity, and sufficiently low viscosity, to travel down the flanks of a volcano, or cross open

Volcanoes

ground. The expression is used both for active flows during emplacement, and for the resulting landform. In the course of a long-lived eruption, a lava flow field may develop by the superposition of many individual flow units. The current eruption of K¯ılauea, Hawai’i, has yielded over 2 km3 of lava and built up a flow field around 100 km2 in area since it began in 1983 (a mean eruption intensity of around 3 m3 s−1 ). Where lava enters the sea, as is the case in Hawai’i, a lava delta or bench may build seawards—though these are often unstable features. As active lava interacts with sea water on the shore in front of the lava bench, steam explosions can construct ephemeral littoral cones. Active lava flows radiate prodigious quantities of heat near the vent such that they rapidly form a surface crust. This can thicken sufficiently to insulate the core of the flow from thermal losses, lowering the rate of viscosity increase, and thereby promoting a longer travel distance. Mafic flows quite often crust over completely, with lava continuing to flow in tunnels, which can grow in cross-section by thermal erosion of the walls. When the supply of lava at the vent ceases, the last slug of lava may drain down-slope leaving an empty conduit or lava tube. On K¯ılauea, much of the lava flow between the Pu’u’ O’o vent, and the coastline where lava pours into the sea (a distance of over 10 km), takes place in a tunnel network, with only sporadic breakouts at the surface. Lava flows can result in dramatic destruction of property (Figure 15.2b), and loss of agricultural land and forest by burial and the action of fires.

Debris Avalanches, Debris Flows, and Lahars Because of their slopes and construction from sometimes poorly consolidated materials (rubbly lavas, loose ash), and through the weakening action of acidic groundwaters, volcanoes are prone to gravitational collapses. These may be triggered by magmatic activity within the edifice, seismic events, or heavy rainfalls. Even small volcanic landslides can be devastating in populated areas, and they can occur long after volcanic activity has ceased. This was the case with the landslide triggered by sustained and heavy rainfall at Casita volcano, Nicaragua, in 1998 (Kerle et al. 2003). In such cases, the landslides typically evolve into mudflows (see below). The major hazards posed by the various kinds of volcanic avalanche and flow phenomena are physical injuries related to burial and property damage, and drowning. Debris avalanches are fast moving, gravity-driven currents of partially or fully water-saturated debris—in

439

this context, primarily volcanic—that are not confined by an established channel. If the moving debris is watersaturated, and enters drainage channels, it is termed a debris flow, and if it consists of a significant fraction of clay-sized particles, it is called a mudflow (or lahar). Such flows can pick up water and debris along the way, while sedimenting out their coarser, denser materials, thereby transforming into clay- and water-rich hyperconcentrated flows. Lahars are a recurrent hazard at many volcanoes worldwide, and have claimed many thousands of lives. They form in a variety of ways, including the rapid melting of snow and ice by hot pyroclastic material, intense rainfall on loose volcanic deposits, and breakout of crater lakes, as well as being a consequence of debris avalanches. They can travel at speeds of 50 km hr−1 or more and travel many tens of kilometres. The lahar arising from the eruption of Nevado del Ruíz Colombia in 1985 covered more than 60 km and resulted in the deaths of an estimated 22,800 people.

Gas and Aerosol Emissions Volcanic gases and particles emitted into the atmosphere are ultimately deposited at the Earth’s surface, where they may have impacts on terrestrial and aquatic ecosystems, agriculture, infrastructure, and human health. The chemical and physical form in which they are deposited, and the spatial and temporal distribution of deposition, are strongly controlled by atmospheric chemistry and the transport of the volcanic plume. Various components of volcanic emissions (including acid species and heavy metals) can affect vegetation, and can have both harmful and beneficial effects. The detrimental effects are generally either mediated through acidification of soils (by dry or wet deposition) or by direct fumigation of foliage (e.g. respiration of acid gases through stomata). Chronic exposures to SO2 concentrations of a few 10s or 100s ppb are sufficient to affect plant ecosystems, decrease agricultural productivity, and cause foliar damage. However, other gas species (e.g., HF and HCl) can be important, as well as the extent of soil acidification due to wet and dry deposition (e.g. Delmelle et al. 2003). Several volcanic volatile species are harmful on contact with the skin, if taken into the lungs, or ingested (Figure 15.3; Hansell and Oppenheimer 2004). Few primary studies have been conducted into health effects of volcanic gases—those that exist are limited in terms of exposure assessment, so the true extent of health effects from volcanic gases is unclear. Most research to date relates to CO2 , H2 S, and SO2 exposures. Sulfur species

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Clive Oppenheimer and David Pyle

Fig. 15.3. Fumarolic and diffuse soil emissions on Vulcano (Italy) pose a health hazard. Needless to say, this sign does not deter thousands of visitors from peering into the ‘smoke holes’ every summer (photo: Clive Oppenheimer).

(SO2 and H2 S gas, sulphate aerosol) can affect respiratory and cardiovascular health in humans (ibid.). Air quality in Hawai’i is reportedly affected by ‘vog’ (volcanic fog) associated with SO2 and sulfate aerosol from Kilauea’s plume, and ‘laze’ (lava haze), composed of HClrich droplets formed when active lavas enter the sea. More recent attention has focused on the abundance of very fine (sub-micron) and very acidic (pH about 1) aerosol emitted from volcanoes (A. G. Allen et al. 2002; R. S. Martin et al. 2008), which is likely to have human health consequences. Accumulations of H2 S in volcanic and geothermal areas, including faulty geothermal heating systems, have resulted in fatalities from asphyxiation. Communities in some geothermal areas are exposed chronically to low H2 S levels, notably in Rotorua, New Zealand, with possible impacts on the nervous system, and on respiratory and cardiovascular health.

Emissions of CO2 can also accumulate dangerously in low-lying areas and have resulted in deaths due to asphyxiation. Several cases have been reported at Furnas (Azores), Vulcano, Lazio, and Alban Hills (in Italy; see below), Cosigüina (Nicaragua), and Mammoth Mountain (California). Dissolved CO2 may also accumulate in lake water in volcanic areas. Sudden displacement of the water may release a cloud of CO2 able to flow under gravity, suffocating people and animals in its path, as occurred at Lakes Nyos and Monoun in Cameroon in the 1980s (Sigurdsson et al. 1987). A further volatile species encountered in volcanic and geothermal areas is radon (e.g. Avino et al. 1999). Many studies have considered an association between radon exposure and lung cancer in humans (e.g. Bowie and Bowie 1991; Avino et al. 1999; J. Pearce and Boyle 2005) raising some concern over the potential exposure in volcanic areas.

Volcanoes

441

TABLE 15.3. Twentieth-century record of deaths, injuries, and other impacts of volcanic activity in the Mediterranean Volcano

Phenomena

Date

Deaths

Injuries

Homeless

TOTAL

(inc. evacuees) Vesuvius Vesuvius Stromboli Santorini Etna Vesuvius Etna Stromboli Vesuvius Campi Flegrei Etna Etna Etna Campi Flegrei Etna Stromboli Etna Stromboli Stromboli Stromboli Stromboli Stromboli

T T(213)/L(3)/G(2) T T L L P T(24)/L(2)/I(1) S L T L S S T T T T T T T

10/03/1905 06/04/1906 22/05/1919 11/08/1925 07/11/1928 03/06/1929 02/08/1929 11/09/1930 00/03/1944 02/03/1970 04/08/1979 12/09/1979 17/03/1981 04/10/1983 16/10/1984 24/07/1986 17/04/1987 16/10/1993 01/06/1996 22/08/1996 04/09/1996 23/08/1999

1 218–700 4

1 300 20

100,000 10 houses many 4,300–5,000 60 houses

2 4–>6 22* 3 9

20 36,000 250 23–4 many houses

1 1 2

15+

250 10 years. This is consistent with eruption durations being controlled by the amount of magma that has accumulated in a shallow chamber between eruptions. (b) The heights of lava domes as a function of time elapsed during an eruption are shown for the 1866 and 1939 Kameni eruptions (grey circles) and compared to the domes of Mt. St Helens and St Vincent. The Kameni domes grew with approximately the same time dependence as the domes of St Helens and St Vincent, consistent with a model of steady inflation of a dome with a stiff outer shell. Since dome height is related to the duration of the eruption, the sizes of domes expected in future eruptions of the Kamenis are predictable (modified from Pyle and Elliott 2006).

violence that apparently included the ejection of large ballistic blocks, followed, over the next three days, by the emission of considerable quantities of noxious gases as well as of ash and steam. Contemporary reports suggest that the gases alone accounted for the deaths of tens of people by suffocation, as well as of countless animals. The climax of the eruption was accompanied by a great wave, usually interpreted as a tsunami, which

The three islands of Kos, Yali, and Nisyros form a cluster of recent volcanic activity in the Dodecanese, close to the Bodrum Peninsula of Turkey (Figure 15.1). Kos, the second largest of the Dodecanese islands, comprises a metamorphic basement, overlain by Mesozoic limestones and flysch. Volcanic activity in the Miocene led to the emplacement of extrusive ignimbrites and intrusive granitoids, and was followed by a period of erosion during which much of western Kos was planed to sea level (Higgins and Higgins 1996). A subsequent period of volcanism began about 3 Myr ago, and resulted in the emplacement of a couple of large rhyolite domes on the south-western Peninsula, and a major pumice dome. This phase culminated in caldera collapse and rhyolite extrusion at around 0.55 Myr ago. The most significant, and recent, activity on Kos was the eruption of the Kos Plateau Tuff at 161 kyr BP (S. R. Allen 2001; Dufek and Bergantz 2007). This eruption, which has been very precisely dated by single crystal laser fusion and Ar-Ar techniques (P. E. Smith et al. 1996), coated much of Kos, and many of the surrounding islands, with pumice fallout and pyroclastic current deposits (Figure 15.4). The discovery of Kos Plateau Tuff deposits on the Turkish mainland led to considerable interest in the question of whether pyroclastic density currents had travelled across the sea (S. R. Allen and Cas 2001). New work on the structure of the sea floor and the age of the marine sediments in

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Clive Oppenheimer and David Pyle

the region suggests, however, that the shelf between Kos and Turkey was much shallower at the time of the eruption, and it looks less likely that the flows travelled across any substantial bodies of water (Pe-Piper et al. 2005). One curious feature of the Kos volcanic sequences is that the deposits tend to be very silica rich (rhyolitic), and the eruptions are widely spaced in time. This contrasts with the much more typical subduction volcano of Santorini, which shows both a greater diversity of melt compositions and a more systematic pattern of repeated eruption. The Kos Plateau Tuff is actually both highly crystalline and silica rich, leading to the suggestion that it may either represent a partial melt of a young granite body or have formed by remobilization of a partially consolidated magma body (Keller 1969). This process, of remobilization of previously stagnated magma bodies, appears to be common in continental margin settings, and may explain the similarities of a number of large- to very largevolume ignimbrite-forming events, including the cataclysmic eruptions of Toba (about 74 kyr ago; Gardner et al. 2002; Oppenheimer 2002) and the Fish Canyon Tuff (around 25 Myr ago; Bachmann et al. 2002). As yet, though, the processes that control when remobilization may occur remain obscure, and our understanding of the potential for future volcanic activity on Kos is limited. The small, 2 km-long, volcanic island of Yali lies south of Kos in the direction of Nisyros. This enigmatic island is thought to be a portion of a partially submerged caldera, and comprises spectacular Late Pleistocene spherulitic obsidian flows to the north, and a Quaternary sequence of pumiceous rhyolites to the south. Isotopic and trace elemental studies show that the Yali rhyolites are distinct from the rhyodacites of Nisyros (Buettner et al. 2005). The >150 m-thick pumice breccia sequence exposed on the south of the island is believed to be the uplifted remnant of a tephra rampart built by a series of submarine and phreatomagmatic eruptions (S. R. Allen and McPhie 2000). Yali has certainly been the site of Holocene activity, since the youngest pumice deposits overlie a sequence of stratified paleosols, one of which contains Neolithic pottery (Keller 1980). Today, the pumice deposits are extensively mined for a variety of commercial uses, but the obsidian, by virtue of its spherulitic nature, was never of a sufficiently high quality for use in blade manufacture (Higgins and Higgins 1996). Hydrothermal activity continues offshore from Yali, as well as in other areas between Kos and Nisyros (Varnavas and Cronan 1991), and the possibil-

ity of future volcanic activity in this region cannot be ruled out. Nisyros is currently the Aegean volcano that is showing the most unrest (Gottsmann et al. 2007). It is a textbook symmetrical stratocone rising sharply out of the sea. The subaerial part forms a small island (with an area of about 42 km2 ), dominated by a central caldera. It comprises a complex of interstratified andesitic-torhyodacitic tephra deposits, probably emplaced between around 60 and 35 kyr ago, which are partially overlain by a younger sequence of large, rhyodacite domes, which now dominate the western portion of the caldera. Historic activity has been mainly phreatomagmatic, producing the extensive fine ash deposits and explosion craters of the Lakki plains that form the eastern portion of the caldera. Nisyros has been extensively studied in the past, but the absence of such excellent and continuous exposure as exists, for example, on Santorini, has made the details of the volcanic history of the island much harder to unravel. Di Paola (1974) produced the first extensive map and descriptions of the island, while the pyroclastic record and tephra dispersal were investigated by Limburg and Varekamp (1991) and Hardiman (1999). Vougioukalakis (1993) has published the most detailed map and stratigraphy to date, and the petrology and isotope geochemistry of the sequences have been extensively described, most recently by Francalanci et al. (1995) and Buettner et al. (2005). It is clear from these studies that recycling of magmatic products— whether melts, crystalline residues from previous melts, or older crustal fragments—has been an important factor in the evolution of Nisyros magmas. This, again, makes the products intrinsically interesting to volcanologists and petrologists, but provides little concrete information for those wishing to understand the future of the magmatic system. In terms of the pyroclastic history of the island, there is still a puzzle over the origin of the caldera: despite many efforts to search for a marine tephra associated with the two major Late Pleistocene silicic explosive eruptions of Nisyros, it is only recently that a candidate Nisyros ash has been recognized in distant, north Aegean, lake and marine cores (Margari et al. 2007; Aksu et al. 2008), and there is scant evidence that either eruption was of sufficiently large magnitude to account for the size of the caldera (Hardiman 1999). Perhaps, instead, the caldera grew incrementally during episodes of both explosive and effusive activity. Of all the volcanoes in the present Aegean arc, Nisyros is the only one that shows any evidence for a present-

Volcanoes

day hydrothermal system on land (Marini et al. 1993; Chiodini et al. 2002). This is a cause for considerable concern, since isotopic studies reveal clear evidence for a magmatic signature in the currently degassing fluids (Brombach et al. 2003; Shimizu et al. 2005). Indeed, the recent seismic (1995–8) and ground deformation crises on Nisyros were accompanied by major gas geochemical changes (1997–2001), which all point towards the prospect of a renewal of the phreatic explosive activity that occurred there in the late nineteenth century (Papadopoulos et al. 1998; Sachpazi et al. 2002; Caliro et al. 2005). The onset of such activity might be challenging to forecast, although the likely focus is clearly identified as the area close to the Lofos dome on the Lakki plain (Caliro et al. 2005). The impact on the island and its tourist economy could be substantial.

Volcanic Risk Management in the Mediterranean One of the main aims of research undertaken to assess and quantify volcanic risks is of course to implement the findings in management practice and policy. Probabilistic risk assessment for current and potential future volcanic hazards is a relatively new endeavour but it is increasingly being used to inform risk management and decision-making. Such assessment makes use of information from a variety of sources, including studies of the rock record associated with past volcanism, data from monitoring networks, theoretical models, and information on the distribution and vulnerability of people or property and assets at risk. Only recently have rigorous, evidence-based approaches emerged in volcanic risk assessment (Aspinall et al. 2003; Baxter et al. 2008a ).

Characterizing, Monitoring, and Modelling Volcanic Hazards One of the first steps in volcanic hazard assessment is characterization of the past history of activity at the volcano in question, in order to build up frequencymagnitude curves for various phenomena, and thereby a statistical basis for long-range forecasting. For instance, a number of studies have attempted to characterize the volcanic hazard presented by the Campi Flegrei, including papers by Lirer et al. (2001b) and Alberico et al. (2002). One of the problems of understanding hazard in such systems where many vents are found is identifying the most likely future vent locations (e.g., see

457

Alberico et al. 2008). Based on the frequency-magnitude relationships for past eruptions, Orsi et al. (2004) suggested that the most probable ‘maximum expected event’ is a medium-magnitude explosive eruption, most likely to be focused on the north-eastern or western sectors of the caldera. It will probably alternate between magmatic and phreatomagmatic phases with the generation of tephra fallout and pyroclastic density currents. They published hazard maps indicating vent location probability, thickness of tephra fall deposits, and areas liable to inundation by pyroclastic density currents. More recently, Mastrolorenzo et al. (2008) produced probabilistic models of tephra fall for a range of Campi Flegrei eruption scenarios and showed that the hazard to buildings is high not only within the caldera but also in Naples due to the prevailing westerly winds. Similarly, Lirer et al. (2001a ) have generated longrange hazard maps for Somma-Vesuvio based on distributions of pyroclastic deposits from the main explosive events that occurred over the last 8 kyr. For tephra fallout hazard, they considered loads exceeding 300 kg m−2 as destructive, and used the relationship between load and frequency to assess relative hazards. They concluded that a total area of around 1500 km2 —which is home to nearly 2 million people—could be affected to varying degrees by the immediate impacts of a future eruption. While such approaches may be statistically robust, they need to be evaluated in the context of improved understanding of the evolution of the magmatic system that feeds the volcano (e.g. Scaillet et al. 2008). For medium- to short-range forecasting of volcanic activity, reliance shifts from the rock record to instrumental monitoring of earthquakes, ground displacements, gas emissions, and so on. Italy may justly lay claim to the establishment of the world’s first volcano observatory, which was constructed on the slopes of Vesuvio in 1841. Today, the modern Osservatorio Vesuviano (Figure 15.15a) is operated by the Istituto Nazionale di Geofisica e Vulcanologia (INGV; ). Other INGV offices include the former Istituto Internazionale per la Vulcanologia founded in Catania in 1969 and now known as INGV ‘Catania section’ (), and a centre for geochemisty based in the University of Palermo (; websites accessed 17 November 2008). In addition to fulfilling routine monitoring responsibilities, these centres and associated university-based groups are engaged in some of the most innovative volcanological research being undertaken today (Figure 15.15b). Many of the capabilities for

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Etna alone are highlighted in a recent publication in the American Geophysical Union’s Geophysical Monograph series (Bonaccorso et al. 2004). Routine monitoring approaches include: r Seismology, which offers an important basis for

r

r

r r

r

r

Fig. 15.15. Volcano monitoring and crisis response. (a) Operations room at the Osservatorio Vesuviano. A bank of monitors displays real-time seismological and other data streamed in from field instruments. (b) Prototype diode laser-based spectrometer for field measurement of volcanic gases being tested on the high temperature fumaroles of Gran Cratere on Vulcano. The optics for the instrument are placed directly over the hot vents with spectra recorded on a laptop in the foreground. This innovative project was one of several exploring new techniques for volcano monitoring funded by the Italian Gruppo Nazionale per la Vulcanologia (GNV). (c) Earth barriers were rapidly constructed in response to the 2001 eruption of Mt. Etna, here to deflect lava flows away from the tourist village (seen from the air in Figure 15.2b) (photos: Clive Oppenheimer).

early detection of volcanic unrest, for location of magma bodies, and for recognizing patterns related to changes in eruptive style (for example, see the work on Stromboli by Jaquet and Carniel 2003, and Langer and Falsaperla 2003). Alaprone et al. (2007) report the success of an automated alert system for tephra fallout based on seismological analysis. Geodetic survey to detect and quantify ground deformation related to magma migration below the surface. For example, both ground-based interferometric spaceborne radar methods are applied at Campi Flegrei (e.g. Lanari et al. 2004). Fluid geochemistry, which can provide evidence for influence of magma degassing in surface volcanic phenomena (e.g. Burton et al. 2003). Thermal surveys using infrared imaging devices (e.g. Calvari et al. 2005). Measurement of microgravity changes related to magmatic and hydrothermal processes (e.g. Gottsmann et al. 2003). Petrological measurements, including characterization of lava geochemistry to infer magma sources and processes (e.g. Corsaro and Miraglia 2005). Doppler radar observations of energetic eruption plumes (e.g. Donnadieu et al. 2005).

Measurements are applied in various modelling approaches. For example, during flank eruptions of Mt. Etna, lava flow hazard is modelled operationally with increasing sophistication (e.g. Barca et al. 2004; Damiani et al. 2006). Indeed there is a considerable body of research that addresses the modelling of various kinds of volcanic flow phenomena (see e.g. Sheridan 2005). Research has also been carried out on the particular behaviour of pyroclastic density currents in both the ancient (Gurioli et al. 2005) and modern (Spence et al. 2004) built environments around Vesuvius. Geographical Information Systems are also finding an increasing role in risk management (e.g. Pareschi et al. 2000a ; Renschler 2005). The specific issues of aviation risk arising from volcanic activity are addressed by the International Airways Volcano Watch (IAVW), which was established by the International Civil Aviation Organization (ICAO) in the late 1980s. The IAVW consists of

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several regional Volcanic Ash Advisory Centres (VAACs), which are responsible for monitoring ash clouds and delivery of warnings to the aviation community. The Mediterranean region is covered by a VAAC in Grenoble, France.

Reducing Risk Science-based evaluation of risks alone is insufficient if the end goal is to protect vulnerable populations. It is now widely acknowledged that systematic development and application of policies, strategies, and practices is essential, and that social and economic policies designed to reduce vulnerability are crucial (see Francis and Oppenheimer 2004; Chapter 18). Communities, and the officials responsible for their protection, need to appreciate not only the nature of risks, but also the uncertainties in the science underpinning the forecasting of volcanic activity and the effectiveness of actions that may be taken to prevent or mitigate the risks. Practicalities of risk management plans also need to be thought through in advance. For example, for successful evacuation, plans need to encompass communications, transport, lodging, medical care, and protection of assets. As a consequence, land use planning, public awareness programmes, education and training are increasingly being undertaken as part of volcanic risk management, for example in the communities threatened by Vesuvius. An acute problem in Italy that has contributed to the exposure of large numbers of people to volcanic hazards is the widespread illegal building known as abusivismo, which has plagued the country for half a century. This phenomenon is strongly manifested on the flanks of Vesuvius, where it has contributed very substantially to the risk from future volcanic activity, as at Campi Flegrei. The establishment in 1995 of the Vesuvius National Park, within a tight legal framework, at last seems to have limited further urban sprawl on the volcano. More recently, dramatic measures have been adopted by the Campania Regional Government aimed at achieving a gradual population decrease in the highest risk area around Somma-Vesuvio. These include cash incentives for families to relocate away from their homes in the zone considered most at risk. It is hoped that up to 100,000 people can be enticed out of this danger zone within fifteen years. In the comuni threatened by Vesuvius, public awareness is also being raised through exhibitions, and via school education. The latter is seen as a particularly effective means of outreach. A relatively small number of teachers introducing information on hazards and

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risks into their curricula can collectively reach a very large number of pupils, who in turn will discuss what they have done at school with their families, in theory spreading positive messages about risk reduction. In the Neapolitan area the public awareness programme could do much to prepare the population in advance of any future activity of Vesuvius. An important component of risk management is the development of emergency plans to be implemented in the event of volcanic crisis. In Italy, the Dipartimento della Protezione Civile has drawn up plans for both Somma–Vesuvio and Campi Flegrei. Recent efforts have focused on evaluating and understanding community preparedness, risk perception, and vulnerability in districts threatened by Vesuvius, including surveys of awareness of the emergency plan developed for the area (e.g. Barberi et al. 2008; Carlino et al. 2008; Solana et al. 2008). While these and other studies have found evidence for a degree of hazard awareness and preparedness in the at-risk population, they have also indicated substantial scope for improvement. The Aegean has experienced far less volcanic activity than Italy in the historic period and it is not surprising therefore that volcanology, in general, is far less prolific in Greece in contrast to work on seismicity and the earthquake hazard (Chapter 16). One consequence is that planning for future volcanic emergencies in Greece appears to be rudimentary; some shortcomings have been highlighted recently for Santorini by DomineyHowes and Minos-Minopoulos (2004). Although techniques have yet to be devised that can prevent volcanic eruptions (and are frankly hard to envisage), numerous engineering solutions are available to modify hazards. Of particular note is the construction of earth barriers for lava flow control on Etna (Barberi and Carapezza 2004). Flank eruptions of the volcano are usually accompanied by remarkable efforts by the civil protection authorities to deflect lava flows away from valuable assets such as the tourist villages on the upper part of the mountain (Figure 15.15c).

Conclusions Millions of people live at risk from volcanism in the Mediterranean area. As we have seen, the hazards presented are extremely varied, and potentially devastating at the regional scale. Reflecting, in part, these high levels of risk, some of the most sophisticated and robust techniques for mapping, monitoring, and modelling volcanoes and their associated hazards and risks have been developed through research on Mediterranean

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volcanoes, notably as a result of major European research projects (e.g., EXPLORIS; Baxter et al. 2008b). Indeed, Etna and Vesuvius, along with Kilauea and Mount St. Helens are among the most researched volcanoes on the planet (and in the solar system for that matter). Many methods that were merely experimental a few years ago are now routinely applied to monitoring of the Italian volcanoes. Santorini, on the other hand, provides one of the most fascinating insights into disasters and human ecology in the ancient world. In the historic period, the largest and deadliest event in the region is the 1631 eruption of Vesuvio. But the ancient and prehistoric records point to even greater and more intense volcanic activity, such as the Late Pleistocene Y5 event, whose recurrence could prove catastrophic. Aside from the threat of very large magnitude eruptions (Mason et al. 2004), given the coastal and island setting of many of the Mediterranean volcanoes, one major issue that needs to be addressed further is the extent to which future eruptions or large scale flank failure could trigger a regional tsunami (Chapter 17). At Santorini, for example, eruptions of both Colombos Reef, and the caldera-collapse accompanying the Bronze Age Minoan eruption have been strongly implicated in the formation of tsunamis (e.g. Cita and Aloisi 2000; McCoy and Heiken 2000; Pareschi et al. 2006). Although the search for related tsunami deposits on the southern shores of Santorini, and the northern coast of Crete has proved equivocal to date (e.g. compare DomineyHowes 2002 and Minoura et al. 2000), it remains unclear whether this reflects a lack of preservation, less intense tsunamis than popularly imagined, or, and perhaps most likely, uncertain criteria for distinguishing the products of a tsunami event from a storm surge or turbidity current triggered by a different process. More recently, Bruins et al. (2008) have identified tsunami deposits at Palaikastro on the eastern coast of Crete that suggested wave heights of 9 m. Whatever the case for the Minoan eruption, we need only consider the welldocumented 1883 eruption of Krakatau to appreciate the potential future scenarios.

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Orsi, G., Civetta, L., Del Gaudio, C., De Vita, S., Di Vito, R. A., Isaia, R., Petrazzuoli, S. M., Ricciardi, G. P., and Ricco, C. (1999), Short-term ground deformations and seismicity in the resurgent Campi Flegrei caldera (Italy): an example of active block-resurgence in a densely populated area. Journal of Volcanology and Geothermal Research 91: 415–51. Di Vito, M. A., and Isaia, R. (2004), Volcanic hazard assessment at the restless Campi Flegrei caldera. Bulletin of Volcanology 66: 514–30. Palladino, D. M. and Simei, S. (2002), Three types of pyroclastic currents and their deposits; examples from the Vulsini volcanoes, Italy. Journal of Volcanology and Geothermal Research 116: 97–118. (2005), Eruptive dynamics and caldera collapse during the Onano eruption, Vulsini, Italy. Bulletin of Volcanology 67: 423–40. Palumbo, A. (1997), Chaos hides and generates order; an application to forecasting the next eruption of Vesuvius. Journal of Volcanology and Geothermal Research 79: 139–48. (1998), Long-term forecasting of the extreme eruptions of Etna. Journal of Volcanology and Geothermal Research 83: 167–72. Papadopoulos, G. A., Sachpazi, M., Panopoulous, G., and Stavrakakis, G. (1998), The volcanoseismic crisis of 1996– 97 in Nisyros, SE Aegean Sea, Greece. Terra Nova 10: 151–4. Pareschi, M. T., Favalli, M., Giannini, F., Sulpizio, R., Zanchetta, G., and Santacroce, R. (2000a ), May 5, 1998, debris flows in circum-Vesuvian areas (southern Italy): Insights for hazard assessment. Geology 28: 639–42. Cavarra, L., Favalli, M., Giannini, F., and Meriggi, A. (2000b), GIS and volcanic risk management. Natural Hazards 21: 361–79. Ranci, M., Valenza, M., and Graziani, G. (2001), Atmospheric dispersion of volcanic CO2 at Vulcano Island. Journal of Volcanology and Geothermal Research 108: 219–35. Santacroce, R., Sulpizio, R., and Zanchetta, G. (2002), Volcaniclastic debris flows in the Clanio Valley (Campania, Italy): insights for the assessment of hazard potential. Geomorphology 43: 219–31. Favalli, M., and Boschi, E. (2006), Impact of the Minoan tsunami of Santorini: Simulated scenarios in the eastern Mediterranean. Geophysical Research Letters 33: L18607, doi: 10.1029/2006GL027205. Patella, D. and Mauriello, P. (1999), The geophysical contribution to the safeguard of historical sites in active volcanic areas; the Vesuvius case-history. Journal of Applied Geophysics 41: 241–58. Paterne, M., Guichard, F., and Labeyrie, J. (1988), Explosive activity of the south Italian volcanoes during the past 80,000 years as determined by marine tephrochronology. Journal of Volcanology and Geothermal Research 34: 153–72. Pearce, J. and Boyle, P. (2005), Examining the relationship between lung cancer and radon in small areas across scotland. Health and Place 11: 275–82. Pearce, N. J. G., Westgate, J. A., Preece, S. J., Eastwood, W. J., and Perkins, W. T. (2004), Identification of Aniakchak (Alaska) tephra in Greenland ice core challenges the 1645 BC date for Minoan eruption of Santorini. Geochemistry Geophysics Geosystems 5: Q03005.

Pègues, Abbé (1842), Histoire et phénomènes du volcan et des îles volcaniques de Santorin: suivis d’un coup d’oeil sur l’état moral et religieux de la Grèce moderne. Imprimerie royale, Paris. Pe-Piper, G. and Piper, D. J. W. (2002), Igneous rocks of Greece: The anatomy of an orogen. Beiträge zur regionalen Geologie der Erde 30: 1–573. Perissoratis, C. (2005), Neotectonics and the Kos Plateau Tuff eruption of 161 ka, South Aegean arc. Journal of Volcanology and Geothermal Research 139: 315–38. Petrazzuoli, S. M. and Zuccaro, G. (2004), Structural resistance of reinforced concrete buildings under pyroclastic flows: a study of the Vesuvian area. Journal of Volcanology and Geothermal Research 133: 353–67. Pizzino, L., Galli, G., Mancini, C., Quattrocchi, F., and Scarlato, P. (2002), Natural gas hazard (CO2 , 222 Rn) within a quiescent volcanic region and its relations with tectonics; the case of the Ciampino-Marino area, Alban Hills Volcano, Italy. Natural Hazards 27: 257–87. Poli, S., Chiesa, S., Gillot, P.-Y., Guichard, F., and Vezzoli, L. (1989), Time dimension in the geochemical approach and hazard estimates of a volcanic area: The isle of Ischia case (Italy). Journal of Volcanology and Geothermal Research 36: 327–35. Principe, C., Tanguy, J. C., Arrighi, S., Paiotti, A., Le Goff, M., and Zoppi, U. (2004), Chronology of Vesuvius’ activity from A.D. 79 to 1631 based on archeomagnetism of lavas and historical sources. Bulletin of Volcanology 66: 703–24. Pyle, D. M. (1990), New volume estimates for the Minoan eruption, in D. Hardy, J. Keller, V. P. Galanopoulos, N. C. Flemming, and T. H. Druitt (eds.), Thera and the Aegean World III. Thera Foundation, London, ii. 113–21. and Elliott, J. R. (2006), Quantitative morphology, recent evolution and future activity of the Kameni islands volcano, Santorini, Greece. Geosphere 2: 253–68. Ricketts, G. D., Margari, V., van Andel, T. H., Sinitsyn, A. A., Praslov, N. D., and Lisitsyn, S. (2006), Wide dispersal and deposition of distal tephra during the Pleistocene ‘Campanian Ignimbrite/Y5’ eruption, Italy. Quaternary Science Reviews 25/21–2: 2713–28. Ramsey, C. B., Manning, S. W., and Galimberti, M. (2004), Dating the volcanic eruption at Thera. Radiocarbon 46: 325–44. Renschler, C. S. (2005), Scales and uncertainties in using models and GIS for volcano hazard prediction. Journal of Volcanology and Geothermal Research 139: 73–87. Rice, A. (2000), Rollover in volcanic crater lakes: a possible cause for Lake Nyos type disasters. Journal of Volcanology and Geothermal Research 97: 233–9. Rinaldi, M. and Venuti, M. C. (2003), The submarine eruption of the Bombarda volcano, Milos Island, Cyclades, Greece. Bulletin of Volcanology 65: 282–93. Rosi, M., Vezzoli, L., Aleotti, P., and De Censi, M. (1996), Interaction between caldera collapse and eruptive dynamics during the Campanian Ignimbrite eruption, Phlegraean Fields, Italy. Bulletin of Volcanology 57: 541–54. Sachpazi, M., Kontoes, C., Voulgaris, N., Laigle, M., Vougioukalakis, G., Sikioti, O., Stavrakakis, G., Baskoutas, J., Kalogeras, J., and Lepine, J. C. (2002), Seismological and SAR signature of unrest at Nisyros caldera, Greece. Journal of Volcanology and Geothermal Research 116: 19–33.

Volcanoes Salvi, F., Scandone, R., and Palma, C. (2006), Statistical analysis of the historical activity of Mount Etna, aimed at the evaluation of volcanic hazard. Journal of Volcanology and Geothermal Research 15: 159–68. Scaillet, B., Pichavant, M., and Cioni, R. (2008), Upward migration of Vesuvius magma chamber over the past 20,000 years. Nature 455: 216–19. Scandone, R., Arganese, G., and Galdi, F. (1993), The evaluation of volcanic risk in the Vesuvian area (in Mount Vesuvius). Journal of Volcanology and Geothermal Research 58: 263–71. Scarth, A. and Tanguy, J.-C. (2001), Volcanoes of Europe. Oxford University Press, Oxford. Schmid, M., Halbwachs, M., Wehrli, B., and Wuest, A. (2005), Weak mixing in lake kivu: New insights indicate increasing risk of uncontrolled gas eruption. Geochemistry Geophysics Geosystems 6: Q07009, doi: 10.1029/2004GC000892. Shackleton, N. J., Fairbanks, R. G., Chiu, T.-C., and Parrenin, F. (2004), Absolute calibration of the Greenland time scale: implications for Antarctic time scales and for ‰14 C. Quaternary Science Reviews 23: 1513–22. Sheridan, M. (ed.) (2005), Modeling and Simulation of Geophysical Mass Flows. Journal of Volcanology and Geothermal Research 3/1–2: 1–146. Shimizu, A., Sumino, H., Nagao, K., Notsu, K., and Mitropoulos, P. (2005), Variation in noble gas isotopic composition of gas samples from the Aegean arc, Greece. Journal of Volcanology and Geothermal Research 140: 321–39. Siebert, L. and Simkin, T. (2002), Volcanoes of the World: An Illustrated Catalog of Holocene Volcanoes and their Eruptions. Smithsonian Institution, Global Volcanism Program Digital Information Series, GVP-3, , accessed 30 October 2008. Sigurdsson, H., Devine, J. D., Tchoua, F. M., Presser, T. C., Pringle, M. K. W., and Evans, W. C. (1987), Origin of the lethal gas burst from Lake Monoun, Cameroun. Journal of Volcanology and Geothermal Research 31: 1–16. Simkin, T. and Siebert, L. (1994), Volcanoes of the World. Geoscience, Tucson. and Blong, R. (2001), Disasters: volcano fatalities— lessons from the historical record. Science 291: 255. Sinitsyn, A. A. (2003), A Palaeolithic ‘Pompeii’ at Kostenki, Russia. Antiquity 77: 9–14. Smith, P. A. and Cronan, D. S. (1983), The geochemistry of metalliferous sediments and waters associated with shallow submarine hydrothermal activity (Santorini, Aegean Sea). Chemical Geology 39: 241–62. Smith, P. E., York, D., Chen, Y., and Evensen, N. M. (1996), Single crystal Ar-40-Ar-39 dating of a late Quaternary paroxysm on Kos, Greece: Concordance of terrestrial and marine ages. Geophysical Research Letters 23: 3047–50. Solana, M. C., Kilburn, C. R. J., and Rolandi, G. (2008), Communicating eruption and hazard forecasts on Vesuvius, Southern Italy. Journal of Volcanology and Geothermal Research 172: 308–14. Sparks, R. S. J. and Wilson, C. J. N. (1990), The Minoan deposits: a review of their characteristics and interpretation, in D. Hardy, J. Keller, V. P. Galanopoulos, N. C. Flemming, and T. H. Druitt (eds.), Thera and the Aegean World III. Thera Foundation, London, ii. 89–99. Spence, R. J. S., Baxter, P. J., and Zuccaro, G. (2004), Building vulnerability and human casualty estimation

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for a pyroclastic flow: A model and its application to Vesuvius. Journal of Volcanology and Geothermal Research 133: 321–43. Stewart, A. L. and McPhie, J. (2003), Internal structure and emplacement of an Upper Pliocene dacite cryptodome, Milos Island, Greece. Journal of Volcanology and Geothermal Research 124: 129–48. Stiros, S. C. and Chasapis, A. (2003), Geodetic monitoring of the Santorini (Thera) volcano. Survey Review 37: 84–8. Stothers, R. B. and Rampino, M. R. (1983), Volcanic eruptions in the Mediterranean before AD 630 from written and archaeological sources. Journal of Geophysical Research 88: 6357–71. Tanguy, J. C., Ribiere, C., Scarth, A., and Tjetjep, W. S. (1998), Victims from volcanic eruptions: a revised database. Bulletin of Volcanology 60: 137–44. Tibaldi, A. and Vezzoli, L. (2004), A new type of volcano flank failure: The resurgent caldera sector collapse, Ischia, Italy. Geophysical Research Letters 31: L14605, doi: 10.1029/2004GL020419. Tinti, S., Bortolucci, E., and Armigliato, A. (1999), Numerical simulation of the landslide-induced tsunami of 1988 on Vulcano Island, Italy. Bulletin of Volcanology 61: 121–37. Pagnoni, G., and Zaniboni, F. (2003), Tsunami generation in Stromboli Island and impact on the south-east Tyrrhenian coasts. Natural Hazards and Earth System Sciences 3: 299–309. Todesco, M., Neri, A., Esposti Ongaro, T., Papale, P., Macedonio, G., Santacroce, R., and Longo, A. (2002), Pyroclastic flow hazard assessment at Vesuvius (Italy) by using numerical modeling; I, Large-scale dynamics. Bulletin of Volcanology 64: 155–77. Ton-That, T., Singer, B., and Paterne, M. (2001), 40 Ar/39 Ar dating of latest Pleistocene (41 ka) marine tephra in the Mediterranean Sea: implications for global climate records. Earth and Planetary Science Letters 184: 645–58. Toyos, G., Oppenheimer, C., Sulpizio, R., Zanchetta, G., and Zuccaro, G. (2003), Modelling building vulnerability to debris flows in the Sarno area, Southern Italy, in D. Rickenmann and C. L. Chen (eds.), Debris-flow Hazards Mitigation: Mechanics, Prediction, and Assessment. Millpress, Rotterdam, 1209–20. van Hinsbergen, D. J. J., Snel, E., Garstman, S. A., Marunteanu, M., Langereis, C. G., Wortel, M. J. R., and Meulenkamp, J. E. (2004), Vertical motions in the Aegean volcanic arc: evidence for rapid subsidence preceding volcanic activity on Milos and Aegina. Marine Geology 209: 329–45. Varnavas, S. P. and Cronan, D. S. (1991), Hydrothermal metallogenic processes off the islands of Nisiros and Kos in the Hellenic Volcanic Arc. Marine Geology 99: 109–33. Vougioukalakis, G. (1993), Volcanic stratigraphy and evolution of Nisyros Island. Bulletin of the Geological Society of Greece 28: 239–58. Witham, C. S. (2005), Volcanic disasters and incidents: A new database. Journal of Volcanology and Geothermal Research 148: 191–233. Wulf, S., Kraml, M., Kuhn, T., Schwarz, M., Inthorn, M., Keller, J., Kuscu, I., and Halbach, P. (2002), Marine tephra from the Cape Riva eruption (22 ka) of Santorini in the Sea of Marmara. Marine Geology 183: 131–41.

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Zanchetta, G., Sulpizio, R., Pareschi. M. T., Leoni, F. M., And Santacroce, R. (2004), Characteristics of May 5–6, 1998 volcaniclastic debris flows in the Sarno area (Campania, southern Italy): relationships to structural damage and hazard zonation. Journal of Volcanology and Geothermal Research 133: 377–93.

Zollo, A., Maercklin, N., Vassallo, M., Dello Iacono, D., Virieux, J., and Gasparini, P. (2008), Seismic reflections reveal a massive melt layer feeding Campi Flegrei caldera. Geophysical Research Letters 35: L12306, doi: 10.1029/ 2008GL034242.

This chapter should be cited as follows Oppenheimer, C. and Pyle, D. M. (2009), Volcanoes, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 435–468.

16

Earthquakes Stathis Stiros

Introduction Earthquakes have played a major role in the evolution of the Mediterranean landscape. They are the most important geohazard in the region and huge sums are invested annually in seismic monitoring, hazard zoning, and earthquake prediction, and in the design of earthquake-resistant buildings and infrastructure. Large earthquakes of magnitude >7.0 have been recorded across the region and the archaeological record shows that earthquakes have posed a major hazard to human settlements for thousands of years (Ambraseys 1971; Shaw et al. 2008; Bottari et al. 2009; Figure 16.1 and Table 16.1). The study of Mediterranean seismicity started about 2,400 years ago when the first earthquake catalogue was compiled in ancient Greece (Papazachos and Papazachou 1997; Guidoboni et al. 1994). This key development predated, by several centuries, the construction of the first seismograph in China (Bullen and Bolt 1985). Since these early developments a great deal of research has been carried out to improve our understanding of earthquakes and associated hazards in the Mediterranean region and to provide protection from them. Earthquake resistant buildings—such as houses with timber bracing—were introduced in Asia Minor in the seventeenth century (Kirikov 1992; Simopoulos 1984; Stiros 1995) and the first strict anti-seismic construction regulations were implemented on the island of Levkas, Greece, in the nineteenth century under British Rule (Stiros 1995). The first ‘modern’, regional-scale earthquake maps and catalogues were compiled as early as the middle of the nineteenth century (Mallet 1858). Despite this progress, the death toll from Mediterranean earthquakes is still high and earthquakes in the region continue to surprise geoscientists. For example, the diffuse pattern of seismicity that is especially char-

acteristic of the eastern Mediterranean (Figure 16.2) is not easily reconciled with existing plate tectonic models, and many faults that are believed to demarcate plate boundaries (such as the Jordan Rift) are currently quiescent (Figure 16.3). Similarly, the 1995 Grevena-Kozani earthquake was a surprise for scientists, for it hit the heart of what was believed to be an aseismic region in northern Greece (Stiros 1998a). Furthermore, key aspects of the geodynamic background of the Mediterranean region remain a matter of debate. This chapter has three main aims. First, to outline the geodynamic setting for Mediterranean seismicity in the context of evolving ideas about plate tectonic theory. Second, to outline the methods that have been developed to study Mediterranean earthquakes and to describe their characteristics. Third, to assess the impacts of Mediterranean earthquakes from both a geomorphological and a geohazard management perspective.

The Geodynamic Background The advent of international global seismographic networks in the mid-twentieth century led to the discovery that most earthquakes are confined to narrow zones associated with lithosphere plate boundaries. The band of seismicity separating the Eurasian and African plates is clear in the Atlantic, but becomes less clear to the east as it traverses the Mediterranean to the east of Algeria (Figures 16.2 and 16.3). It becomes less clear in terms of both the distribution of epicentres and the focal mechanisms of earthquakes (McKenzie 1978; Kiratzi and Papazachos 1995). Plate boundaries can be identified in Turkey (e.g. the North and East Anatolian Faults), but they are not so clearly defined in areas such as Italy, Greece, and western Turkey, where the convergence velocities between Eurasia and Africa

470

Stathis Stiros

(b)

(a)

Fig. 16.1. (a) Ruins of buildings in the village of Verneuges, Provence, that were destroyed by the 1909 Lambesc earthquake (see Baroux et al. 2003). (b) A tombstone from the ancient Greek town of Nikomedia, in the epicentral area of the 1999 North Anatolian Fault earthquakes, Turkey, commemorating the death of two young boys and their teacher after the seismic collapse of their house in AD 120. This tombstone is in the Louvre Museum in Paris (Inscription code CIG–3293) (after Robert 1978).

TABLE 16.1. Some major or otherwise noteworthy earthquakes in the Mediterranean Date

Site

2003

Algerian coast

1999

Athens, Greece

1999

Kocaeli and Duzce earthquakes of the North Anatolian Fault, Turkey

1995

Kozani-Grevena, northern Greece

1981

Corinth, Greece

1980

Al Asnam, Algeria

1980

Irpinia (Campania-Basilicata), Italy Amorgos Island, Central Aegean

1956

1935

Coastal Libya

Effects and important features A major (M = 6.8) earthquake along the Algerian coast which produced the first known case of coastal uplift there. It caused disruption of underwater communication cables and Internet connections across a wide area. A relatively small (M = 5.9), but highly destructive earthquake (over 140 casualties) that triggered a warming of the political climate between Turkey and Greece. The surprising effect of this moderate earthquake was the very high accelerations, possibly locally exceeding 1g. Two major (M = 7.4, 7.2) and destructive (about 20,000 people killed) strike-slip faulting earthquakes along the North Anatolian Fault. The last shocks of a twentieth-century cluster of earthquakes, testifying to progressive rupturing of this fault. They also triggered a warming of the political climate between Turkey and Greece. A surprise M = 6.5 earthquake that revealed that no aseismic zones exist within a region of distributed seismicity—only earthquakes with longer recurrence intervals. A seismic sequence associated with surface ruptures, which allowed the first modelling of normal earthquakes based on a combination of geomorphological observations of seismic and long-term uplift and subsidence, and synthetic (modelled) seismograms. A destructive reverse faulting earthquake associated with surface rupture. Uplift blocked the flow of a river and formed a temporary lake. A major (M = 6.9) destructive earthquake with a high death toll (>3000); the first earthquake in the northern Mediterranean west of the Balkans associated with surface faulting. A magnitude 7.4 earthquake associated with a 20-m high tsunami. This was the largest event in the central Aegean region during the twentieth century and occurred in a region that had formerly been considered aseismic. The death toll was 53, but it would have been much higher if it had occurred 50 years later in the modern era of densely populated coastal areas. A magnitude 7+ earthquake in the apparently aseismic Libya which caused no damage or casualties, as it hit desert area with low population density.

Reference Meghraoui et al. (2004)

Gazetas et al. (2002)

Toksoz et al. (1979); Barka (1996); Stein et al. (1997); Erdik et al. (2003) Stiros (1998a)

Jackson et al. (1982)

Meghraoui et al. (1988) Westaway and Jackson (1984); Pantosti et al. (1993) Ambraseys (1960); Papadopoulos and Pavlides (1992)

Ambraseys (1984)

Earthquakes

471

TABLE 16.1. cont. Date

Site

1927

Damia, Jordan River Valley

1926

Rhodes, Aegean Sea

1909

Provence (France)

1908

Messina, Italy

1894

Locris, central Greece

1825

Levkas Island, Ionian Sea, Greece

1688

Izmir (Smyrni) Turkey

1202

Jordan Valley, Bekaa Valley (Middle East) Antioch, Syria

AD

526

AD

365

Crete, Alexandria, and the eastern Mediterranean

c.AD 250 373

BC

427/6

464

BC

Mavra Litharia, Gulf of Corinth, Greece Helike, central Greece

Locris, central Greece

Sparta, Greece

BC

c.500

BC

Megiddo, Israel

c.550

BC

Possibly Syros Island, central Aegean Jericho

c.1020

BC

c.1500

BC

Kea Island, central Aegean Sea

Effects and important features A magnitude 6 earthquake associated with a landslide that blocked the flow of the Jordan River for several days. This provides a physical explanation for the crossing of the Jordan River by the Jews in c.1020 BC, as described in biblical texts. A magnitude 7.4+ earthquake on Rhodes that led to destruction in Crete. Its effects inspired Sir Arthur Evans to assign the destruction in the palace of Knossos he was excavating to an earthquake. An earthquake of about magnitude 6 that led to deaths and major damage in several towns, mostly because of low-strength buildings. It represents the most damaging seismic event in France for centuries, and probably the first case of an earthquake in which the ‘topography effect’ was noticed, i.e. the amplification of seismic acceleration in the vicinity of sites with high topography gradient. A devastating earthquake of magnitude 7.5 which killed over 60,000 people; about 40% of the total population of the town of Messina and caused up to 70 cm of coastal subsidence. A single span (>2 km) suspension bridge is due to be constructed in the epicentral area of this earthquake. A destructive seismic sequence which inspired the first known pseudo-scientific (unsuccessful) prediction of future earthquakes. A very destructive earthquake. In its aftermath the whole of Levkas town was reconstructed using strict regulations and brilliant anti-seismic techniques that were introduced as a building code under British rule. After this destructive earthquake, stone-built houses were replaced by timber-reinforced houses that responded well to subsequent earthquakes. A large (magnitude 7.5+), destructive earthquake felt at distances of >2,000 km away. It was associated with a tsunami (Chapter 17). Near-total destruction from seismic shocks and a conflagration with possibly 250,000 people killed. In AD 587/8 another destructive earthquake killed another perhaps 60,000. A giant earthquake that became a legend; its anniversary, ‘the Day of Horror’ was commemorated for centuries in Egypt. This earthquake produced a tsunami that destroyed Egyptian towns. It has similarities with the December 2004 Asian tsunami. This earthquake, probably a series of earthquakes of magnitude over 8.5, produced >9 m uplift in Crete and nearly total destruction of towns. The 365 earthquake belongs to a cluster of major earthquakes across the eastern Mediterranean known as the ‘Early Byzantine tectonic Paroxysm’ sensu Pirazzoli (1986). Several metres of coastal uplift destroyed the harbour of the flourishing town of Aigeira, which was subsequently all but abandoned. A rather common earthquake that, several centuries later, became a legend and its effects (destruction, tsunami, subsidence of a coastal strip of land) exaggerated and explained as punishment for a sin city. One of the very first earthquakes known to have been associated with coastal subsidence and changes in the flow of springs. The only surviving entry from the first catalogue of earthquakes in Greece more than 2,000 years ago. The first case of an earthquake that triggered a revolution: seismic destruction of Sparta triggered a revolt of slaves and subordinate nations, and the Spartan regime survived thanks to the support of the rival state of Athens, which was interested in preserving social stability. A major earthquake that marked the end of a c.5,000 year old town that had survived numerous earlier earthquakes. Memories of this event are possibly imprinted in biblical texts including the Armageddon earthquake in the Revelation of St John. The first reported ‘prediction’ of an earthquake based on non-scientific reasoning. This earthquake is claimed to be responsible for collapse of the walls of Jericho, permitting its capture by the Jewish army. An example of an earthquake with wider physical effects and historical impacts: due to seismic coastal subsidence, a spring chamber was contaminated by saline water and the site was then abandoned.

Reference Nur and Ron (1996)

Ambraseys and Adams (1998); Stiros (1995) Baroux et al. (2003); and unpublished data from the author

Valensise and Pantosti (1992)

Unpublished data from the author Papazachos and Papazachou (1997); Stiros (1995) Stiros (1995) Ambraseys and Melville (1988) Guidoboni et al. (1994)

Guidoboni et al. (1994); Pirazzoli et al. (1996); Pirazzoli (1986); Stiros (2001); Stiros and Papageorgiou (2001) Stiros (1998b) This study

Guidoboni et al. (1994) Papazachos and Papazachou (1997) Thucidides, 1.101.2; Stiros (1996)

Nur and Ron (1996)

Guidoboni et al. (1994) Nur and Ron (1996) Stiros (2005)

472

Stathis Stiros

60°N

45°N

30°N

15°N

0°

30°W

15°W

0°°

15°E

30°E

45°E

Fig. 16.2. Epicentres of shallow earthquakes in the Mediterranean reported by the USGS during 1961–83 (modified from Jackson and McKenzie (1988) with additions). Epicentres are confined to a narrow zone in the Atlantic and the Red Sea, but are distributed over wide zones especially in Greece and western Turkey, making the identification of plate boundaries difficult. Areas free of earthquake epicentres exist between regions of distributed deformation including the Adriatic and the central-southern Aegean. The arrow points to the previously assumed aseismic area of western Macedonia (Greece) which was hit by the 1995 Kozani-Grevena earthquake.

become gradually higher (DeMets et al. 1990). This has led to intensive debate as to whether simple or more complicated styles of plate boundary exist between Eurasia and Africa as the Mediterranean has progressively narrowed during the final stages of the now extinct Tethys Ocean (Chapter 1). On the basis of the distribution of the focal mechanisms of earthquakes, Anderson and Jackson (1987) have shown that the Adriatic cannot be regarded as an African promontory, as was previously believed, while more recent Global Positioning System (GPS) data provide some evidence to suggest that the boundary between the African and Eurasian plates crosses Sicily (Hollenstein et al. 2003).

To address these plate boundary problems in the Aegean and western Turkey, McKenzie (1972) introduced the hypothesis of micro-plates. More recent studies, however, favour the hypothesis of continuous, distributed deformation across broad regions (Figure 16.4; Jackson 1993b) and of incipient rifting. The latter is a tendency for the formation of major faultcontrolled grabens and these are expressed through a new generation of small east–west trending faults at a strike favourable to accommodate strain from the North Anatolian Fault to the Ionian Sea which replace the older north–west trending faults (Hatzfeld 1999). This strain pattern is indeed confirmed by strike-slip

Earthquakes

473

Epicentres 1900-1965 Shallow

8>M>7 7>M>6 6>M>5

Intermediate Epicentral regions I – XVII Centuries

250 km

Fig. 16.3. Epicentres of instrumentally recorded earthquakes (M > 5) 1900–65, and areas affected by historical earthquakes in the eastern Mediterranean (modified from Ambraseys 1971). Note the absence of twentieth-century seismicity along some of the major plate boundaries such as the Dead-Sea Fault.

earthquakes along the North Anatolian Fault and by recent GPS data showing a westward motion of the Anatolian plate (McClusky et al. 2000). The geodynamic setting is complex, however, with many unresolved issues. Three main reasons for these uncertainties can be identified: 1. The radius of the Mediterranean arcs (e.g. the Aegean arc) is too short and the length of the subducting plate slab too small (about 200 km long) relative to those of typical island arcs of the Pacific and with variable dip (Hatzfeld et al. 1993). Detailed seismic profiling has indicated a spatial correlation between zones of middle crust mobilized by flow and an upper crust thinned by oro-

genic extension and expressed near the surface as an exhumed metamorphic core complex (Sachpazi et al. 1997). 2. The focal mechanisms of earthquakes used to derive plate movements in this area are rather limited in number. The older ones in particular are not well constrained and their pattern is not clear. For example, in the central part of the Hellenic arc they show east–west instead of the expected north–south compression (Taymaz et al. 1990). Many models assume that the eastern edge of the Hellenic arc is essentially a strike-slip margin (Figure 16.4) (Jackson 1993a), but the late Holocene and longer-term uplift history of Rhodes indicates an important thrusting component, as

474

Stathis Stiros

Eurasia

NAF

F

EA

Africa

DSF

Arabia

250 km

Fig. 16.4. Plate boundaries in the eastern Mediterranean (modified from Jackson 1993a). The hashed zone marks the area of distributed deformation between the North Anatolian Fault (NAF) and the Ionian Sea. The arrows indicate dominant plate movements. The East Anatolian Fault (EAF) and the Dead Sea Fault (DSF) are also shown. The solid triangles mark the Hellenic (Aegean) arc subduction zone between the African and Eurasian plates.

seen in seismic reflection profiles (Kontogianni et al. 2002; Kiratzi and Papazachos 1995). 3. Matters become rather more complicated in some key areas where deformation appears not to be controlled by seismicity and a deficit exists between deformation derived from instrumentally recorded seismicity and deformation deduced from tectonic and geodetic models. For example, in peninsular Italy or central Greece, no such deficit exists (Hunstad and England 1999; Ambraseys and Jackson 1990). However, in the Hellenic arc, where seismicity rates are high, instrumentally recorded earthquakes account for only a small part of the tectonic strain released. This indicates either dominant aseismic deformation (Jackson 1993a; Jackson and McKenzie 1988; Jenny et al. 2004), consistent with geodetic evidence (Billiris et al. 1991; Stiros 1993), or the occurrence of strong earthquakes with long recurrence intervals (Stiros 1998a; Stiros and Drakos 2006), or perhaps even earthquake swarms (Pirazzoli 1986; Nur and Cline 2000; Stiros 2001; Ambraseys 2004). Thessaly in central Greece, for example, was practically aseismic for centuries, but was marked by seven earthquakes of magnitude > 6 between 1930 and 1980 (see Papazachos and Papazachou 1997; Caputo and Helly 2005). Similarly, the Dead Sea Fault, and its northern extension up to the East Anatolian Fault, representing a plate boundary, appears almost aseismic in the instrumental seismicity records (Figures 16.2 and 16.3). They

were, however, reactivated during a cluster of strong earthquakes (magnitude up to 7.5+) between AD 1114 and 1202 (Ambraseys and Melville 1988; Ambraseys 2004). Similarly, while the Ionian Sea and the Peloponnese currently together represent one of the most seismically active parts of the Mediterranean, the scarcity of both historical seismicity data (Guidoboni et al. 1994; Papazachos and Papazachou 1997; Albini 2004) and of coastal palaeoseismicity data (Pirazzoli et al. 1996), suggests that is was rather aseismic in antiquity (Figure 16.5).

Methods for Studying Mediterranean Earthquakes Seismological (sensu stricto) methods focus on the identification of epicentres and the depth of earthquakes based on the analysis of recordings from various monitoring stations and the compilation of fault-plane solutions using various techniques (see Jackson 2001). Instrumental seismicity, however, scarcely covers the last hundred years or so and the database has to be supplemented by other approaches. Indeed, a key feature of the Mediterranean is that the seismic history of most areas can be reconstructed for thousands of years because of the rich historical records and excellent preservation of ancient remains (Ambraseys 1971; Stiros and Jones 1996; Bottari et al. 2009). This serves to counter some of the problems discussed above that have been met in the study of Mediterranean seismicity because of the tectonic complexity of the region.

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(a) Seismicity 1963–1988

(b) Maximum observed intensities 1700–1981

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(c) Intense seismicity according to Strabo (1st Century AD)

IX VIII VII

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Fig. 16.5. Seismicity across the Peloponnese during three periods: (a) Instrumentally derived seismicity (modified from Jackson 1993b). (b) Maximum observed earthquake intensities 1700–1981 using the MKS scale (based on IGME 1989). (c) Seismically active areas according to the first-century AD writer Strabo. There is a poor correlation between the seismicity pattern in the three periods and this may reflect time-dependent seismicity in this region.

It is not coincidental that the subdisciplines of Historical Seismology and Archaeoseismology were developed in the Mediterranean region (Jones and Stiros 2000). Historical Seismology is based on 2,500 years of historical records that yield information on seismic events

over this period. Key earthquake parameters such as epicentre location and earthquake magnitude and intensity can be estimated (Ambraseys 1971; Guidoboni et al. 1994; Stucchi and Camassi 1996). Archaeoseismology is based on ancient remains bearing traces of ancient earthquakes (Table 16.2) which typically

TABLE 16.2. A list of indicative criteria for the identification of earthquakes from archaeological data. Since various phenomena can produce similar effects to earthquakes, their identification requires that at least one of these criteria are satisfied and preferably in more than one building or site, and also that other possible causes of destruction can be excluded 1. 2. 3. 4.

Ancient buildings, walls etc. offset by seismic surface faults. Human skeletons buried under the debris of fallen buildings. Abrupt geomorphological changes that may be associated with the destruction and/or abandonment of buildings and sites. Characteristic structural damage and failure of built structures: r Displaced drums of dry masonry columns r Opened vertical joints and horizontal displacement of walls in dry masonry walls r Diagonal cracks in rigid walls r Triangular sections missing in the corners of masonry buildings r Cracks at the base or top of masonry columns and piers r Inclined or sub-vertical cracks in the upper parts of rigid arches, vaults, and domes, or their partial collapse along these cracks r Vertically displaced keystones in dry masonry arches and vaults r Several parallel fallen columns r Several fallen columns with their drums in a domino-style (‘slices of salami’) arrangement r Built structures deformed by horizontal forces so that rectangular forms are transformed to parallelograms. 5. The destruction and rapid reconstruction of sites, with the introduction of what can be regarded as ‘anti-seismic’ building construction techniques, but with no change in their overall culture. 6. Well-dated building collapse events that correlate with known earthquakes from historical sources. 7. Damage to, or destruction of, isolated buildings or entire sites and for which an earthquake provides the only feasible explanation. Sources: Modified from Stiros (1996); Karcz and Kafri (1978).

476

(a (a)

Stathis Stiros

(b)

(c)

Fig. 16.6. Examples of earthquake damage to ancient buildings: (a) Offsets of the drums of the marble columns of the Hephaisteion (Thesion) Temple near the Acropolis in Athens, the best-preserved ancient Greek temple. These offsets, assigned to an earthquake by the early seismologist Kritikos, led Galanopoulos (1956) to suggest that Athens was not aseismic, as was hitherto believed, but in the past had been shaken by destructive earthquakes. This idea was confirmed by the 1999 Athens earthquake, which, despite its relatively small (140) death toll. Such column drum offsets are a result of the oscillation of the columns by strong horizontal seismic forces (Sinopoli 1991; Stiros 1996). (b) Columns of a Roman villa in Susita, Galilee, close to the Dead Sea Transform fault, that were probably toppled during the AD 365 earthquake (after Nur and Ron 1996). (c) The parallel walls of Triolo Temple—a Greek temple at Selinunte, Sicily—that were damaged by an earthquake in the fourth-century BC (after Bottari 2003).

occurred in the last 3,000 to 4,000 years. This is possible because the occupation of many sites was permanent and because dwellings and other infrastructure (e.g. defensive walls, water channels, etc.) were constructed from stone and mortar, and even from sculpted hard rock, limestone, and marble. Many constructions had a strict geometric plan form so that any deformation associated with earthquakes or other natural effects can be determined (Lanciani 1918; Karcz et al. 1977; Stiros 1996; Hancock and Altunel 1997; Jones and Stiros 2000; Figure 16.6). Recent work by Bottari et al. (2009) based on surveys of ancient monuments in Sicily has shown that this part of the Mediterranean was struck by at least three major earthquakes between 400 BC and AD 600. The events were dated using coins, pottery, and other artefacts. Palaeoseismological studies commonly involve the study of geomorphological and stratigraphical anomalies caused by earthquakes that were associated with surface deformation and faulting. Such studies commonly involve observations from excavated trenches and they have yielded important data on, for example, the nature and extent of seismic slip and, in contexts where the age of the features can be established, recurrence intervals for major earthquakes can be estimated (Figure 16.7) (Meghraoui et al. 1988; Pantosti et al. 1993; Michetti et al. 1996; Chatzipetros et al. 1998). Another advantage for the study of seismicity in the Mediterranean is the presence of long bedrock coasts—frequently in hard carbonate rocks—that pre-

serve signs of seismic crustal uplift (Pirazzoli et al. 1982, 1996) (Figure 16.8; Chapter 13). These rocks can also preserve data on subsidence, occasionally in association with archaeological data (Flemming 1978; Pirazzoli et al. 1982). This situation is aided by the negligible tidal range (usually 100-km-long sector of Crete and this led to the formation of coastal plains more than 500 m wide (Figure 16.12) (Pirazzoli et al. 1982; Kelletat 1991). The AD 365 event also led to a major tsunami that struck much of the eastern Mediterranean (Stiros and Drakos 2006; Shaw et al. 2008; Chapter 17). This event also forced down-cutting in the lower reaches of many of the river channels of western Crete in response to the abrupt change in base level (Figure 16.13). The most recently recorded cases of significant earthquake-related uplift in the Mediterranean region come from the Cephalonia islands, Greece, in the Ionian Sea (Figure 16.8) and the Algerian coast (Meghraoui et al. 2004). The most important example of recent coastal land subsidence is that associated with the 1981 Gulf of Corinth earthquake (Jackson et al. 1982; Aubert et al. 1996). Interestingly, seismic coastal uplift occurs where widely used, large-scale tectonic models predict either crustal extension-associated subsidence (Flemming 1978)—such as in the central part of the Aegean (Pirazzoli et al. 1996; Stiros et al. 2000), or only minor uplift. The water line of the Roman harbour of Aigeira

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Fig. 16.11. (a) A normal fault produced by the 1954 earthquake in Thessaly, central Greece (after Papastamatiou and Mouyiaris 1984). (b) A schematic figure to illustrate reverse faulting, folding, and uplift that dammed the flow of the Cheliff River during the 1980 Al Asnam earthquake in Algeria. The vertical movements associated with this earthquake are shown on the graph below (modified from Stein and Yeats 1989). (c) Railway tracks offset by strike-slip faulting during the 1999 Izmit earthquake in Turkey (photo: A. Barka).

in the Gulf of Corinth (Greece) is presently at the height of 4 m above mean sea level due to one or more earthquakes (Stiros 1998b), while at the nearby rocky shore, marine fossils testify to a spectacular mean average uplift rate of 3 mm per year for the late Holocene (Pirazzoli et al. 2004; Chapter 13). In many cases coastal uplift is followed by further geomorphological changes that can modify the coastal topography. For example, in response to coastal uplift, the lower reaches of streams have incised their channels in order to adjust their gradients (Figure 16.7). In some cases, however, where stream power was relatively small and incision has not kept pace with uplift, dry valleys have formed in permeable lithologies as is the case along the southern coast of the Gulf

of Corinth during the Late Pleistocene (Armijo et al. 1996). At some locations in the Mediterranean, the presence of uplifted raised beach platforms of various ages indicates the persistence of vertical seismic movements and their significance in long-term landscape development. This is the case in western Crete where Holocene, Pleistocene, and Neogene surfaces are present (Pirazzoli et al. 1982; Le Pichon and Angelier 1979). In contrast, recent seismic uplift can be evident in areas where there are no signs of older events, as is the case of the 2003 coastal uplift in Algeria (Meghraoui et al. 2004). There is some evidence that the relationship between earthquakes producing uplift and longterm landscape evolution (derived from elevated marine

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Stathis Stiros

Iraklion

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Falasarna

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Chania

Kastelli

Rethymnon Paleochora

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Isobases marking the uplift of 1550 BP in metres Emerged plains more than 500 metres wide

0

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20 km

Fig. 16.12. Contours of uplift across western Crete resulting from the AD 365 event (modified from Kelletat 1991). See Pirazzoli et al. (1982) for a comparison as these authors present slightly different values of uplift. Emergent coastal plains greater than 500 m wide associated with this earthquake are also shown.

sediments indicating long-term uplift) in the Mediterranean was understood thousands of years ago. For instance, the fifth-century BC poet Pindar described the island of Rhodes as a plot of land emerging from the sea. Rhodes provides a very clear example of how good agreement can be achieved from different sources of evidence. In this case these are: (1) Holocene uplift rates deduced from raised shorelines yielding uplift rates of up to 0.6 mm per year (Pirazzoli et al. 1989); (2) the remains of a c.2,000-year-old harbour that is now several metres above sea level, and (3) exposed marine sediments that are widespread across the island (Kontogianni et al. 2002). Interestingly, Pindar, among other ancient authors, reported that earthquakes were responsible for the opening of a narrow valley, through which water flows to the Aegean and the transformation of a large lake into the present-day Thessaly Plain of central Greece—

the floor of which is composed of lacustrine sediments (Helly 1989). In areas of normal faulting, such effects tend to lower the relief in contrast to the activity of reverse faults, which tend to produce uplift that blocks the flow of rivers and forms lakes. The 1980 earthquake in Al Asnam, Algeria, for example, was the last in a series of thrust-type earthquakes producing uplift and transient or longer-term blocking of the flow of the Chelif River and the formation of a lake (King and Vita-Finzi 1981; Meghraoui et al. 1988).

Secondary Effects In addition to tectonically controlled crustal movements, earthquakes can produce secondary geomorphological effects, such as landslides (Chapter 6), tsunamis (Chapter 17), liquefaction, and the compaction of unconsolidated

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Fig. 16.13. Response of a stream at Sougia, on the coast of south-west Crete, to a c.6 m uplift during the AD 365 earthquake. The upper line of notches to the right and the line marking the top of the dark area to the left indicate sea level prior to the earthquake and the corresponding base level of the stream. The stream bed has adjusted to the post-seismic sea level by incision. Based on the measurements of uplift, it was found that this earthquake was of magnitude >8.5 and was associated with a thrust offshore (Stiros and Drakos 2006). The parallel lines to the right mark the remains of shorelines that were produced by a series (>10) of earthquakes. These led to subsidence (with an average of about 0.3 m for each event) over a period of 2,000 years before the AD 365 earthquake (Pirazzoli et al. 1982 1996; Kelletat 1991). This sequence testifies to a very unusual earthquake cycle (Stiros 1996). (Photo: Stathis Stiros).

sediments. Such secondary effects are mainly due to the influence of seismic waves (i.e. ground shaking) and can have very significant and long-term impacts on local communities.

Landslides Numerous earthquakes are known to have been associated with landslides (Guidoboni et al. 1994; Papazachos and Papazachou 1997; Chapter 6) and these have been shown to disrupt the flow of the River Jordan, for example (Nur and Ron 1996). Earthquake-triggered landslides may increase sediment supply to rivers and cause the burial of parts of towns, as archaeological excavations in various parts of Greece, including Olympia have shown (Stiros 1996, 2001). In Tiryns (near Mycene, southern Greece), in particular, the burial of the lower town by debris seems to have followed the seismic

destruction of a dam about 2,800 years ago (Kilian 1996).

Tsunamis The magnitude 7.4 earthquake that struck Amorgos Island in the central Aegean Sea in 1956 produced a tsunami wave up to 20 m high (Ambraseys 1960). The tsunami associated with the AD 365 earthquake was responsible for the inundation of extensive areas of coastal land in Egypt and the transformation of cultivated land into swamps (Guidoboni et al. 1994; Stiros 2001; Shaw et al. 2008). This process was evident during the 2004 tsunami in south-east Asia. The Syrian coast has been particularly affected by tsunamis produced by earthquake-triggered submarine slumping (Chapter 17).

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Stathis Stiros

Submerged Settlements in Antiquity

Fig. 16.14. An early engraving illustrating the combined effects of ground sliding, compaction, and perhaps liquefaction following the 1783 earthquake in Calabria, Italy (after Michetti 1994).

Compaction and Liquefaction of Unconsolidated Sediments Earthquake shaking can lead to the compaction of unconsolidated sediments sediments (see e.g. Papathanassiou et al. 2005) and this in turn can lead to ground subsidence. Liquefaction is a process whereby unconsolidated surficial sediments lose their cohesion and rigidity and this can have a highly destructive impact on overlying structures (Figure 16.14).

Hydrological Effects Strong ground shaking can change the flow rates from springs and initiate new springs as underground drainage networks become diverted and disconnected or blocked with debris (see Papazachos and Papazachou 1997).

Sinkholes Seismic events may lead to the weakening and collapse of bedrock ceilings above underground voids in karst terrains to form sinkholes. Large boulders in rockshelter and cave sites in the Mediterranean (e.g. Franchthi Cave in southern Greece) have been attributed to seismic shaking (Farrand 2000).

Other Secondary Effects Tectonic deformation may have a very wide number of indirect effects on geomorphological processes. For example, even a relatively small scale (c.1 m) uplift or subsidence at a coastal site may change the direction and magnitude of stream discharge and modify erosion and sedimentation dynamics in the catchment and at the coast.

In some cases there are historical reports of catastrophic earthquake-associated landscape changes and these commonly refer to ancient sites or parts of towns ‘swallowed by the waves’ (see Papazachos and Papazachou 1997; Nikonov 1996). However, many of these claims cannot be accepted as reliable. The best-known case is that of Helike, which along with the nearby town of Boura, lying on a hill, was supposed to have been submerged in the Gulf of Corinth with all its inhabitants during the 373 BC earthquake and tsunami that followed (Marinatos 1960). Such an extraordinary calamity, however, was first mentioned several centuries later in the case of Helike and became a legend. It was not mentioned by any of the contemporary historians, who only provided brief and rather vague reports of an earthquake followed by a small tsunami—events not unusual in the region. This author believes that the legend of the earthquake-related drowning of Helike and Boura, among other similar legends of destruction in the region, should be discarded for the following reasons: 1. Greek towns of this period covered extensive areas (ancient Athens with its hundred demoi, or communities, is the typical example) and no signs of the associated large-scale geomorphological changes that are needed to account for the subsidence of both Helike (located in the coastal plain) and of Boura (located on a hill) have ever been identified. 2. It is very likely that the ‘drowning of Helike’ is a myth that was first propagated (possibly following a low-magnitude earthquake), to disguise its political annihilation by the nearby rival town of Aigion which, miraculously, was not affected by the hypothetical giant earthquake and tsunami, and became the dominant town in the area (Faraklas 1998). 3. There are no natural events or combinations of events (e.g. seismic faulting, landslide, tsunami inundation, liquefaction of loose sediments) that can adequately explain the simultaneous loss of the two towns in their contrasting topographic settings. 4. There is some stratigraphic evidence for buildings and undisturbed occupation levels that predate and postdate the critical period (373 BC) in the region inland (Soter and Katsonopoulou 1999), and this argues strongly against the occurrence of a catastrophic landslide, land subsidence, or any other similar calamity.

Earthquakes

The Historical and Social Significance of Earthquakes in the Mediterranean Earthquakes have often been invoked as a convenient solution to explain discontinuities in an occupation sequence or in the cultural history of various sites and regions and this approach has been criticized by various investigators (Ambraseys 1971; Stiros 1996: Faraklas 1998). This is not to say that, in numerous cases, earthquakes did not play an important role in the history of some sites by, for example, influencing their political role, their architectural and urban (re)modelling, or even their abandonment. The 464 BC earthquake that destroyed Sparta triggered a revolt of slaves and oppressed tribes, and a war that risked destroying the Spartan state. Also, following its demise due to the AD 365 earthquake, Paphos was replaced by Salamis as the capital of Cyprus (Stiros 2001), while after destruction by an earthquake the town of Al Asnam in Algeria was re-established but with a different name. Archaeological excavations have indicated a radical remodelling of some towns in the aftermath of earthquakes (Stiros and Jones 1996). The redesign of Troy prior to the period corresponding to the Trojan War was assigned by Blegen (1963) to the effects of an earthquake. Similarly, the Roman town of Kisamos in Crete that was levelled during the AD 365 earthquake was totally reconstructed as a new, Christian town (Stiros and Papageorgiou 2001). Some modern examples include Salon de Provence in southern France, the centre of which was remodelled after the 1909 earthquake, and some modern Greek earthquakes such as the 1995 Aigion earthquake in the Gulf of Corinth (Koukouvelas and Doutsos 1996). The latter provided an opportunity for the people of this town to replace their traditional twostorey houses made of stone and unbaked bricks even though they formed part of their cultural heritage and were protected by law. There is an important lesson here because they were replaced by modern, multistorey, reinforced-concrete buildings that, in many cases, proved more vulnerable to earthquake destruction than the well-preserved traditional dwellings (Figure 16.15). Finally, there are numerous cases of sites and towns abandoned or relocated following earthquakes. Pella in Macedonia, northern Greece, the birthplace of Alexander the Great, was relocated a few hundred metres away from its original site after a destructive earthquake in around 90 BC (Stiros 1998a). Of the numer-

485

ous archaeological sites and ruins currently bearing names with the prefix ‘Palaio-’, meaning ‘old’, many reflect abandonment after a seismic disaster. Santorini (Thera) Island was largely abandoned after the 1956 earthquakes, until the tourist boom of the 1980s, while after the 1881 Chios Island (central Aegean Sea) earthquakes, the major part of the population left the island, ignoring the orders of the Turkish Government trying to avoid its abandonment. The same situation arose after the 1953 earthquakes in the Ionian Islands. In order to avoid the vacation of the devastated islands, the Greek Government confiscated the fishing boats, thereby condemning many fishermen to starvation. In fact the role of the central government and of other communities has been decisive in the post-earthquake history of towns. During the 464 BC earthquake that devastated Sparta, the Athenians sent an army to support their eternal rivals to prevent a slave revolt and to maintain the oppression of conquered nations. Several ancient towns are known to have survived seismic disasters thanks to financial assistance from the central government or other communities or colonies (Guidoboni 1989). Without such assistance, many settlements effectively disintegrated and ceased to function as organized towns. A good example is Gortyn (the former capital of Crete), Egypt, and Libya during Roman times (Di Vita 1996). In some cases, however, earthquakes were responsible for natural effects that had a catalyzing role in the inhabitation history of sites. For example, around 3,500 years BP, the prehistoric site of Kea Island in the central Aegean was effectively abandoned when a spring that provided the major source of potable water was contaminated by saline water following coastal seismic subsidence (Stiros 2005). Finally, it has been argued that the passage of the Jews to Canaan, as described by the Bible, was in fact made possible because of an earthquakeformed landslide dam across the Jordan River, similar to the dam that formed during an early twentieth century earthquake (Nur and Ron 1996).

Hazard Management: Defence Against Earthquakes In some parts of the world, in the aftermath of a destructive earthquake, many people have little choice but to remain in the affected region because of restrictions imposed by the availability of suitable land and building materials and the nature of the climate. New houses are commonly built in the same locations using the same techniques, even if they have proved susceptible

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Stathis Stiros

(a)

(b)

Fig. 16.15. (a) A collapsed multistorey building that led to several casualties in Kalamata, southern Greece, following the 1986 earthquake. (b) During the same event many traditional buildings were badly damaged but did not collapse, and many of their occupants survived (see Anagnostopoulos et al. 1987). (Photos: Stathis Stiros).

to earthquakes in the past. The settlements that comprise vaulted, unbaked brick houses in the central, arid part of Iran are a good example (see Jackson 2001). In contrast, in most parts of the Mediterranean the availability of land, building materials, and good trading contacts with other regions has allowed people to invent, test, and adopt techniques for anti-seismic construction to reduce the vulnerability of their dwellings. The use of timber framing in houses and iron clamps in the hewn marble blocks of ancient temples are among the most effective anti-seismic protection techniques that have been employed (Kirikov 1992; Stiros 1995). In some cases, it is known that entire towns were rebuilt after earthquakes abandoning the old building styles and introducing anti-seismic construction techniques. Indeed, this was the case in late seventeenth-century Smyrni (Izmir) in western Turkey, after the earthquake of 1688 (Simopoulos 1984). Due to the high death toll in many Mediterranean countries with over 1,000 in Turkey and 19 per year in Greece, on average, throughout the twentieth century (Erdik et al. 2003; A. Anagnostopoulos 2003), seismic protection has become a government priority in most countries. Government-funded research in Greece has focused on understanding earthquakes, including both their history and geography (i.e. the compilation of earthquake catalogues based on instrumental (Macropoulos and Burton 1981), historical (Guidoboni et al. 1994; Ambraseys et al. 1994), or a variety of data (Amiran et al. 1994; Papazachos and Papazachou 1997) and the identification of potential seismic sources (for instance Stein et al. 1997; Valensise and Pantosti

2001). Investment has also been directed towards the compilation of maps of active faults (for instance in Turkey, see Saroglu et al. 1992) of observed seismic intensities (e.g. in Greece: IGME 1989) and of seismic risk, at local, national, and international levels (Erdik et al. 2003; Jimenez et al. 2003). Particular attention has been given to the introduction of regulations for earthquake-resistant buildings and for the development of emergency plans. In Greece an effective anti-seismic protection building code was first introduced on the island of Levkas in the Ionian Sea during the nineteenth century under British Rule (Stiros 1995), while more formal antiseismic building codes were introduced after the 1928 Corinth earthquakes. These were inspired partly by the Japanese codes and were adopted nationally across Greece in 1959. More comprehensive codes to deal with the reinforced concrete buildings that represent the great majority of modern construction in Greece were introduced in 1984 and 1992 (see A. Anagnostopoulos 2003). These codes comprise a map of seismic risk zonation and a list of regulations for the design of earthquakeresistant buildings. Seismic risk zonation is a primary responsibility of seismologists, geologists, and engineers. It is based on our understanding of past earthquakes and incorporates statistical estimates of future earthquakes—mainly estimates of recurrence intervals and expected maximum seismic ground acceleration. While efforts have been made to estimate the seismic risk in the Mediterranean using both time-dependent and time-independent seismicity models (Jimenez et al. 2003;

Earthquakes

Papaioannou and Papazachos 2000), it is difficult to accommodate in these models the time-dependent contrasts in seismic history as exemplified in western Crete, for example. While coastal and archaeological data provide evidence of a magnitude >8.7 earthquake in AD 365 (Pirazzoli et al. 1982; Kelletat 1991; Stiros and Papageorgiou 2001; Stiros and Drakos 2006; Shaw et al. 2008), the seismotectonic map of Greece, based on recorded seismic intensities in the last three hundred years, classifies this area as of low seismic risk (IGME 1989)! Seismic zoning maps are updated from time to time following major earthquakes (e.g Anastasiadis et al. 2001) or the acquisition of new data from palaeoseismic studies. Seismic zoning is not limited to earthquakeprone countries such as Italy and Turkey (Turkish General Directorate of Maps 1972, 1996), but has been adopted for countries such as France, which have much lower levels of seismicity (Desperoux and Godefroy 1986; Figure 16.2). Regulations for the construction of earthquake-resistant buildings (mostly relating to reinforced concrete structures) were first introduced in Greece and later by most Mediterranean countries. Recently there has been a move towards the adoption of unified codes (Eurocodes) that are similar to those in Japan and the United States. However, in most Mediterranean countries such codes are not mandatory and their use is commonly limited to new or remodelled buildings. Much of the existing infrastructure across the Mediterranean region therefore remains vulnerable to the earthquake hazard. It is worth pointing out that seismic risk analysis, including the identification of fault segments with the potential to rupture in a future earthquake—as is the case with the NAF (Toksoz et al. 1979; Barka 1996; Stein et al. 1997)—has little in common with various efforts for ‘short-term’ prediction, which, it can be argued, have largely proved unsuccessful (Mulargia and Geller 2003). Finally, although there have been concerted efforts for international collaboration—especially within the European Union— as far as earthquake research (seismic risk estimation), preparedness (earthquake-resistant infrastructure and emergency planning), and post-event recovery (crisis management and long-term recovery) are concerned (ibid.), there is still much work to do.

References Albini, P. (2004), A survey of past earthquakes in the Eastern Adriatic (14th to early 19th century). Annals of Geophysics 47: 675–704.

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Papastamatiou, D. and Mouyiaris, N. (1986), The Sophades earthquake on 30 April 1954: Field observations by Yannis Papastamatiou [in Greek]. Geology and Geophysics Research Special Issue Athens, 341–62. Papathanassiou, G., Pavlides, S., Christaras, B. and Pitilakis, K. (2005), Liquefaction case histories and empirical relations of earthquake magnitude versus distance from the broader Aegean region. Journal of Geodynamics 40/2–3: 257–78. Papazachos, B., and Papazachou, C. (1997), The Earthquakes of Greece. Zitis, Thessaloniki. Pirazzoli, P. A. (1986), The Early Byzantine tectonic paroxysm. Zeitung fur Geomorphologie, Suppl. NS 62: 31–49. Montaggioni, L. F., Saliege, J., Seconzac, G., Thommeret, Y., and Vergnaud-Grazzini, C. (1989), Crustal block movements from Holocene shorelines: Rhodes island (Greece). Tectonophysics 170: 89–114. Thommeret, J., Thommeret, Y., Laborel, J., and Montagionni, L. (1982), Crustal block movements from Holocene shorelines: Crete and Antikythira (Greece). Tectonophysics 68: 27–43. Laborel, J. and Stiros, S. C. (1996), Earthquake clustering in the Eastern Mediterranean during historical times. Journal of Geophysical Research 101: 6083–97. Stiros, S. C., Fontugne, M., and Arnold, M. (2004), Holocene and Quaternary uplift in the central part of the southern coast of the Corinth Gulf (Greece), Marine Geology 212: 35–44. Robert, L. (1978), Documents d’Asie Mineure. Bulletin de Correspondence Hellenique 102: 395–408. Sachpazi, M., Hirn, A., Nercessian, A., Avedik., F., McBride, J., Loucoyannakis, M., Nicolich., R., and the STREAMERSPROFILES group (1997), A first coincident normal-incidence and wide-angle approach to studying the extending Aegean crust. Tectonophysics 270: 301–12. Saroglu, F., Emre, O., and Kuscu, I. (1992), Active Fault Map of Turkey, 1:1,000,000 Scale, General Directorate of Mineral Research and Exploration (MTA), Ankara. Shaw, B., Ambraseys, N. N., England, P. C., Floyd, M. A., Gorman, G. J., Higham, T. F. G., Jackson, J. A., Nocquet, J. M., Pain, C. C., and Piggott, M. D. (2008), Eastern Mediterranean tectonics and tsunami hazard inferred from the AD 365 earthquake. Nature Geoscience 1: 268–76. Simopoulos, K. (1984), Foreign Travellers in Greece, 1700–1800. 4th edn. Athens, ii [in Greek]. Sinopoli, A. (1991), Dynamic analysis of a stone column excited by a sine wave ground motion. Applied Mechanics Review 44: S246–S255. Soter, S. and Katsonopoulou, D. (1999), Occupation horizons found in the search for the ancient Greek city of Helike. Geoarchaeology: An International Journal 14: 531–63. Stein, R. S. and Yeats, R. S. (1989), Hidden earthquakes. Scientific American 260 (June): 48–57. Barka, A., and Dieterich, J. (1997), Progressive failure on the North Anatolian Fault since 1939 by earthquake stress triggering. Geophysical Journal International 128: 594–604. Stiros, S. C. (1993), Kinematics and deformation of central and southwestern Greece from historical triangulation data and implications for the active tectonics of the Aegean, Tectonophysics 220: 283–300.

(1995), Archaeological evidence of antiseismic constructions in antiquity. Annali di Geofisica 38: 725–36. (1996), Identification of earthquakes from archaeological data: Methodology, criteria and limitations, in Stiros and Jones (1996: 129–52). (1998a), Historical seismicity, palaeoseismicity and seismic risk in Western Macedonia, Northern Greece. Journal of Geodynamics 26: 271–87. (1998b), Archaeological evidence for unusually rapid Holocene uplift rates in an active normal faulting terrain: Roman harbour of Aigeira, Gulf of Corinth, Greece. Geoarchaeology: An International Journal 13: 731–41. (2001), The AD 365 Crete earthquake and possible seismic clustering during the 4–6th centuries AD in the Eastern Mediterranean: a review of historical and archaeological data. Journal of Structural Geology 23: 545–62. (2005), Social and historical impacts of earthquakerelated sea-level changes on ancient (prehistoric to Roman) coastal sites. Zeitschrift für Geomorphologie, Suppl. 137: 79–89. and Drakos, A. (2000), Geodetic constraints on the fault pattern of the 1978, Thessaloniki (Northern Greece), earthquake (Ms = 6.4), Geophysical Journal International 143: 679–88. (2006), A fault-model for the tsunami-associated, magnitude ≥8.5 Eastern Mediterranean, AD 365 earthquake. Zeitschrift für Geomorphologie 146: 125–37. and Jones, R. E. (eds.) (1996), Archaeoseismology, Fitch Laboratory Occasional Paper No. 7. Oxbow, Oxford. and Papageorgiou, S. (2001), Seismicity of Western Crete and the destruction of the town of Kisamos at AD 365: Archaeological evidence. Journal of Seismology 5: 381–97. Pirazzoli, P. A., Laborel, J., and Laborel-Deguen, F. (1994), The 1953 earthquake in Cephalonia (Western Hellenic Arc): coastal uplift and halotectonic faulting. Geophysical Journal International 117: 834–49. Laborel, J., Laborel-Deguen, F., Papageorgiou, S., Evin, J., and Pirazzoli, P. (2000), Seismic coastal uplift in a region of subsidence: Holocene raised shorelines of Samos Island, Aegean Sea, Greece. Marine Geology 170: 41– 58. Stucchi, M. and Camassi, R. (1996), Building up a parametric earthquake catalogue in Europe: the historical background, in D. Giardini (ed.): Proceedings of the NATO Conference on Historic and pre-Historic Earthquakes. Kluwer, Dordrecht. Taymaz, T., Jackson, J., and Westaway, R. (1990), Earthquake mechanisms in the Hellenic Trench near Crete. Geophysical Journal International 102: 695–731. Toksoz, M. N., Shakal, A. F., and Michael, A. J. (1979), Spacetime migration of earthquakes along the North Anatolian Fault Zone and seismic gaps. Pure and Applied Geophysics 117: 1258–70. Turkish General Directorate of Maps (1972), Seismic Zones of Turkey (unpublished). Turkish General Directorate of Maps, Ankara. (1996), Seismic Zones of Turkey, 1:1,800,000 scale. Turkish General Directorate of Maps, Ankara. Valensise, G. and Pantosti, D. (1992), A 125 kyr-long geological record of seismic source repeatability: the Messina Straits (southern Italy), and the 1908 earthquake (Ms 71/2). Terra Nova 4: 472–83.

Earthquakes and Pantosti, D. (2001), The investigation of potential earthquake sources in penisular Italy: A review. Journal of Seismology 5: 287–306. Ward, S. N. and Valensise, G. R. (1989), Fault parameters and slip distribution of the 1915 Avezzano, Italy, earthquake derived form geodetic observations. Bulletin of the Seismological Society of America 79: 690– 710.

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This chapter should be cited as follows Stiros, S. C. (2009), Earthquakes, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 469–491.

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17

Tsunamis Gerassimos Papadopoulos

Introduction According to Imamura (1937: 123), the term tunami or tsunami is a combination of the Japanese word tu (meaning a port) and nami (a long wave), hence long wave in a harbour. He goes on to say that the meaning might also be defined as a seismic sea-wave since most tsunamis are produced by a sudden dip-slip motion along faults during major earthquakes (Figure 17.1). Other submarine or coastal phenomena, however, such as volcanic eruptions, landslides, and gas escapes, are also known to cause tsunamis. According to Van Dorn (1968), ‘tsunami’ is the Japanese name for the gravity wave system formed in the sea following any large-scale, short-duration disturbance of the free surface. Tsunamis fall under the general classification of long waves. The length of the waves is of the order of several tens or hundreds of kilometres and tsunamis usually consist of a series of waves that approach the coast with periods ranging from 5 to 90 minutes (Murty 1977). Some commonly used terms that describe tsunami wave propagation and inundation are illustrated in Figure 17.2. Because of the active lithospheric plate convergence, the Mediterranean area is geodynamically characterized by significant volcanism and high seismicity as discussed in Chapters 15 and 16 respectively. Furthermore, coastal and submarine landslides are quite frequent and this is partly in response to the steep terrain of much of the basin (Papadopoulos et al. 2007a). Tsunamis are among the most remarkable phenomena associated with earthquakes, volcanic eruptions, and landslides in the Mediterranean basin. Until recently, however, it was widely believed that tsunamis either did not occur in the Mediterranean Sea, or they were so rare that they did not pose a threat to coastal communities. Catastrophic tsunamis are more frequent on

Pacific Ocean coasts where both local and transoceanic tsunamis have been documented (Soloviev 1970). In contrast, large tsunami recurrence in the Mediterranean is of the order of several decades and the memory of tsunamis is short-lived. Most people are only aware of the extreme Late Bronge Age tsunami that has been linked to the powerful eruption of Thera volcano in the south Aegean Sea (Marinatos 1939; Chapter 15). Even that wave is commonly not viewed as a significant geophysical event. Indeed, it is often seen as rather an exotic episode—part myth and part fact—given that it happened in prehistoric times and has been linked with the collapse of the Minoan civilization of Crete. These are some of the reasons why, in comparison to other parts of the world, the scientific study of tsunamis in the Mediterranean Sea has been rather neglected until the last few decades. This chapter reviews the tsunami history of the Mediterranean and evaluates the progress that has been made in assessing the tsunami hazard over its several regions. In addition, the prospects for further development of tsunami science and associated risk mitigation technology in the region are outlined.

Tsunami Quantification A parameter that is of particular importance for understanding tsunami-generating mechanisms and assessing tsunami hazard better is wave size, and this can be expressed as either intensity or magnitude. These parameters, however, are difficult to determine— even for more recent events (e.g. Soloviev 1970; Shuto 1993). In Europe, tsunami intensity (k) is traditionally estimated according to the 6-grade SiebergAmbraseys scale (Ambraseys 1962). In this chapter,

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Gerassimos Papadopoulos Outer ridge

Trench

Sanriku

1896

(a)

A

Continent

A

1933

Ocean

Intense inelastic deformation

Continental lithosphere

Oceanic lithosphere

(b) Aleutian Continent

1946 1929

Ocean

Stress concentration Sudden interplate faulting

(c)

Tsunamigenic or aseismic complex faulting

Fig. 17.1. Co-seismic dip-slip motion along faults and tsunamigenesis near deep sea trenches. The diagram on the left (modified from Fukao 1979) shows the sequence (a–c) of inelastic deformation and stress concentration culminating in aseismic faulting. In the Sanriku (north-east Japan) and Aleutian examples, thrust faulting is assumed to be the mechanism responsible for the 1896 and 1946 tsunamigenic earthquakes, whilst normal faulting is assumed for the 1933 and 1929 events (modified from Kanamori 1972).

Maximum run-up r

Wave direction Maximum water levell Tsunami height at shore Tsunami

Sea level at time of tsunami Inundation = maximum horizontal intrusion

Fig. 17.2. A schematic explanation of some of the tsunami terms used in this chapter (modified from IOC 1998).

the reported tsunami intensities are based on this scale, while the Murty-Loomis (1980) definition is adopted for tsunami magnitude (ML). An attempt has been made to recalculate the intensities of some important Mediterranean tsunamis according to a recently introduced 12-grade intensity scale (Papadopoulos and Imamura 2001; Papadopoulos 2003a) and these are shown in Table 17.1.

Major Tsunami Events in the Mediterranean In this section the tsunami history of the Mediterranean Sea is reviewed. Emphasis is given to a selection of key tsunami events and these were chosen because of their significant impact on coastal communities and also because they have been extensively studied by a wide range of geophysical, geological, geomorphological, historical, and archaeological research methods as well as by hydrodynamic numerical modelling techniques. This review also aims to evaluate the advantages and disadvantages of the research methods that have been used, to highlight the progress achieved in this area, and to identify the emerging prospects for further research. The tsunami catalogues used in the following review are based mainly on those of Galanopoulos (1960), Ambraseys (1962), Antonopoulos (1979), Papadopoulos and Chalkis (1984), Tinti and Maramai (1996), Soloviev et al. (2000), Papadopoulos (2001 2003b), Tinti et al. (2004) and Fokaefs

Tsunamis TABLE 17.1. Strong tsunamis of intensity k ≥ 4 reported for the Mediterranean Sea between 426 No

Year

Month

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18 19 20 21 22 23 24 25 26 27 28 29 30 31 32 33 34 35 36 37 38 39 40 41 42 43

426 BC 373 BC 142/4 365 447 551 552 556 749 1169 1202 1303 1343 1365 1389 1402 1481 1609 1612 1627 1650 1693 1741 1748 1759 1766 1773 1783 1817 1823 1856 1866 1866 1866 1867 1908 1944 1948 1956 1963 1979 1999 2002

summer winter

Day

07 01 07 05

21 26 09

01 02 05 08 10 01 03 06 05 04 11 07 10 01 01 05 11 05 05 02 08 03 11 01 02 03 09 12 08 02 07 02 04 08 12

18 04 20 08 18 02 20 03 08 30 11 11 31 25 25 22 06 06 23 05 13 02 02 06 09 28 20 09 09 07 15 17 30

Location

k

Ref

K

Maliakos Gulf W. Gulf of Corinth Rhodes Crete Sea of Marmara Lebanon Maliakos Gulf Kos Levantine coast Messina Straits Syrian coast & Cyprus Crete Sea of Marmara Algiers Chios Gulf of Corinth Rhodes Rhodes Crete Gargano, Italy Thera Eastern Sicily Rhodes W. Gulf of Corinth Akko Sea of Marmara Tangiers Calabria W. Gulf of Corinth North Sicily Chios Albania Kythira Albania Gythion Messina Straits Stromboli Karpathos Cyclades W. Gulf of Corinth Montenegro Sea of Marmara Stromboli

5 5 4 5 4 5 4 4 4 4 4 5 4 4 4 4 4 5 4 4 6 4 5 4 5 4 4 6 4 4 4 4 4 4 4 6 4 4 6 4 4 4 4

P P P P P ** P P ** TM AM P P ** P P ** P P ** P TM P P ** P ** TM P TM P P P P P TM TM P ** P P P *

8 9 7 10 8 8 8 8 7 8 7 10 8 8 6 8 7 8 8 6 10 7 8 9 8 7 7 9 9 8 8 7 6 7 7 10 7 7 9 7 8 6 7

BC

and

AD

495

2002

h (cm)

ML

200

2,000

−1.4 +3.0 +2.3

1,000

900 900 500

−1.8

800

1,300

−0.4

1,500 500

+3.0 −11.0

250 900

Notes: k = tsunami intensity in the 6-grade scale of Sieberg-Ambraseys, K = tsunami intensity in the 12-grade scale of Papadopoulos and Imamura (2001), h = run-up height, ML = Murty-Loomis (1980) tsunami magnitude. * new event. ** revised in this chapter. Sources: AM = Ambraseys (1962), P = Papadopoulos (2001), TM = Tinti and Maramai (1996) and Tinti et al. (2004).

and Papadopoulos (2007). Tsunamigenic zones in the Mediterranean determined from historical data are shown in Figure 17.3a, while Table 17.1 lists the key characteristics of the known strong tsunamis (k ≥ 4). A range of field evidence has been utilized to identify palaeotsunamis and this includes sedimentological, geomorphological, and archaeological data (Figure 17.3b). The following sections review the evidence for tsunami

activity across the Mediterranean region beginning with the Alboran Sea in the west and moving eastwards.

Alboran Sea According to documentary sources, the earthquakes in northern Algeria of 2 January 1365, 6 May 1773, and 21 and 22 August 1856 caused tsunamis of up to

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Gerassimos Papadopoulos

(a)

Potential Low Intermediate High Very high 2

7 3

15

8 5

16 13 12

4 1

14 11

6

18

9

0

1 2 3 4 5 6 7 8 9 10 11 12 13 14 15 16 17 18

10

17

500 km

Alboran Sea Liguria and the Côte d’Azur Tuscany Aeolian Islands Tyrrhenian Calabria Eastern Sicily and Messina Straits Gargano Peninsula East Adriatic Sea Western Hellenic arc Eastern Hellenic arc Cyclades Gulf of Corinth Maliakos Gulf East Aegean Sea North Aegean Sea Sea of Marmara Cyprus Levantine Sea

(b)

Crete Otranto-Leuca

Sarkoy Troy

Itea (3)

Didim Dalaman (3) Thera

Fetiye

Cyprus

Palestine Israel

Archaeological

Geomorphological

Sediment deposits

Pumice

250 km

Fig. 17.3. (a) Tsunamigenic zones in the Mediterranean Sea defined from documentary sources and classified according to their relative tsunami potential as explained in the text. See also Table 17.1. Most of North Africa has not been classified. Here tsunami potential is generally regarded as being low and long-term records are not available. (b) Types of field evidence for the occurrence of past tsunamis.

k 4 (Figure 17.3a). Eyewitness accounts and tide-gauge records verify that similar tsunamis were produced by the strong earthquakes of 9 September 1954, 10 October 1980, and 21 May 2003. In these examples, the location of the earthquakes on land in northern Algeria produced submarine slumping that generated powerful turbidity currents (Soloviev et al. 2000). As a result, atmospheric gravity waves (rissagas) of up to 2-m amplitude and with a frequency of about 10 minutes were observed along the Spanish coast. These waves have also

been observed in southern bays of the Balearic Islands (Monserrat et al. 1991) and also in the Aegean Sea (Papadopoulos 1993a).

Liguria and the Côte d’Azur On 16 October 1979 a submarine slope failure that occurred during the construction of the new Nice airport produced a tsunami wave 3 m high that was observed near Antibes (Figure 17.3a). The near field wave heights

Tsunamis

were successfully simulated by Assier-Rzadkiewicz et al. (2000). The theoretical results were not, however, in complete agreement with the far field observations. This may be explained by the rapid amplitude attenuation of the tsunami due to strong wave dispersion, a common feature of landslide-generated tsunamis (Papadopoulos and Kortekaas 2003; Papadopoulos et al. 2007a). The 20 July 1564 and 23 February 1887 earthquakes (Eva and Rabinovich 1997) triggered tsunamis inundating the coast from Nice to Antibes and from Genoa to Cannes respectively. A modern-day repeat of these intermediate magnitude waves would threaten the densely populated coastal zone of Liguria and the Côte d’Azur.

Tyrrhenian Sea Along the coast of Tuscany in the northern Tyrrhenian Sea (Figure 17.3a), very few events have been reported since the k 4 tsunami of 5 March 1823. On the slopes of Stromboli in the Aeolian Islands (Figures 17.3a and 17.4), volcanic landslides produced tsunamis of k 3 or 4 on 3 July 1916, 22 May 1919, 11 September 1930, 20 August 1944, and 30 December 2002 (Figure 17.5). Further south an extreme event occurred in Tyrrhenian Calabria (Figures 17.3a and 17.4) on 6 February 1783 when a huge earthquake-induced rockfall triggered a k 6 tsunami at the Scilla beach (Tinti and Guidoboni 1988). Inundation heights of 6–9 m were observed and more than 1,500 lives were lost. At Torre del Faro, to the north of Messina, the run-up height was about 6 m and twenty-six people were swept out to sea.

Eastern Sicily and the Messina Straits The 11 January 1693 earthquake in north-east Sicily, that claimed about 70,000 victims, caused a tsunami of k 4. Sea-level oscillations destroyed many boats

Italy Tyrrhenian Sea

South Adriatic Sea

Stromboli 1783 1823

Sicily

1908 Reggio Calabria Messina 1693

1169 100 km

Fig. 17.4. Important tsunamis reported for southern Italy.

497

and ships while flooding was reported from Catania, Augusta, and Messina (Figures 17.3a and 17.4). Tinti et al. (2001) reported hydrodynamical studies of this event and concluded that the faults most likely to be the earthquake source are located in the Scordia-Lentini graben that intercepts the coastline. The 28 December 1908 earthquake is one of the largest ever reported in Italy (Chapter 16). Major towns in southern Italy, including Messina and Reggio Calabria, were completely destroyed with more than 60,000 victims. The earthquake generated a violent tsunami in the Messina Straits (Figures 17.3a and 17.4). This tsunami consisted of at least three large waves that caused many deaths and severe damage to ships, buildings, and property. A tsunami of k 6 was observed along the Calabrian coast at Pellaro and on the Sicilian coast at S. Alessio where wave heights observed were 13 m and 11.7 m respectively. A tsunami with very similar characteristics to the 1908 event is believed to have struck the Messina straits in 1169. A review of the seismological, geological, and geodetic data (Valensise and Pantosti 1992) in association with information on tectonic stress inversion (Neri et al. 2004) indicates that the 1908 earthquake was associated with normal faulting (Chapter 16). However, the seismogenic fault has not yet been identified and this has hindered attempts to develop simulation models of this major Mediterranean tsunami.

Adriatic Sea Along the Adriatic coast of Italy, a tsunami source is related to the seismicity of the Gargano promontory (Figure 17.3a). The destructive earthquake of 30 July 1627, which may have been associated with the Apricena normal fault on land (Patacca and Scandone 2004), caused a k 5 tsunami. Tinti et al. (1995) calculated that large events are expected in this area on average every 228 years. On the eastern side of the Adriatic Sea moderate to strong tsunami events were reported in Albania in 1866 and in Montenegro in 1979 (Figure 17.3a).

The Hellenic Arc The large tsunamigenic earthquake of AD 21 July 365, located off the shore of western Crete (Papadopoulos and Vassilopoulou 2001), is one of the most contentious and debated natural events in Mediterranean Sea history (Chapter 16). The accounts of Marcellinus, Athanasius, and Jerome, which are the closest in time to the event, leave no doubt that a large area was affected since the tsunami propagated to the north-west, west, and south of the Hellenic trench and reached as far as Methoni,

498

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Fig. 17.5. Some of the damage caused by the Stromboli tsunami of 30 December 2002 (photo: Gerassimos Papadopoulos).

Sicily, Alexandria, and Dalmatia (Guidoboni et al. 1994; Ambraseys et al. 1994; Shaw et al. 2008) (Figures 17.3a and 17.6). Coastal uplift of up to 9 m in western Crete and 3 m in Antikythira (Thommeret et al. 1981; Pirazzoli et al. 1992) have been dated to (calibrated ages) AD 341–439 and 265–491, respectively. Therefore, the 365 event probably corresponds to the dramatic uplift event that raised the harbour at Phalasarna in northwest Crete, by c.6.6 m (Chapter 16). In Phalasarna, tsunami deposits attributed to the AD 365 wave were described by Pirazzoli et al. (1992). According to Dominey-Howes et al. (1998), however, there is no biostratigraphical or lithostratigraphical evidence to infer tsunami sedimentary deposition. This may be due to the 6.6 m of uplift which could have taken place only a few minutes before the wave arrived if the uplift in western Crete and the AD 365 tsunami were caused by the same seismotectonic movement. Shaw et al. (2008) have recently argued that this earthquake took place not on the subduction interface beneath Crete, but on a fault with a dip of c.30◦ within the

overriding plate. Their tsunami propagation calculations produced a damaging tsunami wave throughout the eastern Mediterranean with a repeat time of about 5,000 years for such an event. A range of archaeological, stratigraphical, geomorphological, and radiometric data imply that a tsunami struck Phalasarna after a strong earthquake around AD 66 (Pirazzoli et al. 1992; Dominey-Howes et al. 1998). In more recent times, remarkable earthquake tsunamis were also observed between the Peloponnese and Crete on 9 March 1630, 6 February 1866, and 20 September 1867. Documentary sources indicate that the 8 August 1303 earthquake, which ruptured the eastern Hellenic arc between Crete and Rhodes (Figures 17.3a and 17.6), was one of the largest reported for the historical period in the Mediterranean. The most serious destruction was reported from eastern Crete (Guidoboni and Comastri 1997) where a large tsunami struck Iraklion on the north coast. The sea swept violently into the city with such force that it destroyed buildings and killed inhabitants. In Acre, Israel, people were swept away and

Tsunamis

499

TURKEY

GREECE

Methoni

Dalaman

556

Kos

1609 Fig. 17.8

1866 Antikythira

Rhodes

1481 144 1741

South Aegean Sea Karpathos

1612

Phalasarna Iraklion

365 0

Crete

1948 1303

100 km

Fig. 17.6. Important tsunamis along the Hellenic arc. The Cyclades region (boxed area) is shown in greater detail in Figure 17.8.

drowned by a huge wave and in Alexandria the sea destroyed port facilities. In the easternmost Hellenic arc strong earthquake tsunamis occurred in AD 142 or 144 in Rhodes and Kos, AD 556 in Kos, 3 May 1481, April 1609, 31 January 1741, and 23 May 1851 in Rhodes, and 9 February 1948 in Karpathos (Figures 17.3a and 17.6). Trenches in Holocene sediments at Dalaman, south-west Turkey, have revealed three tsunami sand layers (Figure 17.7) attributed to the 1303, 1481, and 1741 tsunamis respectively (Papadopoulos et al. 2004). The 1609 wave, which violently inundated Rhodes, is missing from the Dalaman stratigraphy. This may be due to the fact that it failed either to penetrate this far inland or to leave a tsunami deposit at the location of the trench site.

The Cyclades The Cyclades has a long record of tsunami events and is a key area for tsunami research (Papadopoulos et al. 2007b). The range of tsunami-generating mechanisms and the rich archaeological and sedimentary record of past tsunami activity mean that it is a key natural laboratory for the investigation of tsunami generation and their impacts. The Minoan eruption is one of the most significant because of its size (Volcanic Explosivity Index = 7),

its possible impact on Late Bronze Age (LBA) civilizations, and the distribution of huge amounts of tephra, thereby creating an important marker horizon around the eastern Mediterranean. The eruption history may have included four main phases (Heiken and McCoy 1984) and concluded with the formation of the caldera that dominates the landscape of Thera (Santorini) (Figures 17.3a and 17.8) (Chapter 15). The most intensive eruption phase lasted for about 3–4 days (Sigurdsson et al. 1990). The Thera event has similarities with Krakatau where the collapse of the volcanogenerated tsunamis that rolled against the shores of Java and Sumatra, with heights up to 35 m, leading to more than 36,400 casualties. From archaeological observations on Amnissos in northern Crete, Marinatos (1939) suggested that the Theran tsunami was linked with the demise of the Minoan civilization. In coastal sites, assemblages of usually rounded pumice, often mixed with seashells, have been attributed to the Minoan tsunami. However, several cases are rather problematic. On Anafi Island, pumice layers at altitudes up to 250 m could not be attributed to Minoan tsunami deposition, as Marinos and Melidonis (1971) suggested, because the pumice there is at least 18,000 years old and of air-borne origin (Keller 1978). Pumice found by Marinatos (1939) in Amnissos was linked to the Theran eruption without any analysis. The

500

Gerassimos Papadopoulos

Naxos

Amorgos Ios

1956 Astypalaea

1650 Thera Kamari

Source of Minoan tsunami (17th Century BC)

Anafi

25 km

Fig. 17.8. Important tsunamis in the Cyclades Islands in the southern Aegean Sea.

Fig. 17.7. An excavated section showing palaeotsunami deposits in Dalaman, south-west Turkey. This section was excavated in 1996 and was approximately 230 m from the present shoreline. Three tsunami sediment layers (dark sediments) are present and these correspond to the 1303, 1481, and 1741 tsunamis in the east Hellenic arc (photo: Gerassimos Papadopoulos).

relationship between the Minoan tsunami and pumice deposits found in coastal sites of Cyprus, Israel, and Palestine is also rather speculative (Francaviglia 1990). Minoura et al. (2000) identified Minoan tsunami deposits in Gouves in northern Crete, and in coastal trenches in Didim and Fethye in south-west Turkey. AMS 14 C dating on fossil shells from Didim and on marine gastropod shells from Fethye, indicated that the eruption may have occurred around the second half of the nineteenth century BC and this places it about 200 years earlier than the previous estimates. In Thera, a 3.5-m thick volcaniclastic deposit intercalated with

third and fourth phase deposits of the Late Bronze Age eruption has been interpreted as tephra reworked by the Minoan tsunami (McCoy and Heiken 2000). Dominey-Howes (2004) found no evidence for any Late Bronze Age tsunami at forty-one coastal sites on Crete and Kos. Futher evidence for the Minoan tsunami comes from the marine sedimentary record. Seismic-reflection surveys in topographic lows of the western Mediterranean and Calabrian Ridges have shown a distinct, acoustically transparent, flat-lying layer, nicknamed ‘homogenites’. These deposits occupy the uppermost part of the sediment column (Kastens and Cita 1981) within a stratigraphic unit characterized by an upward fining grain size that implies deposition in a single event controlled by gravitational settling. Kastens and Cita (1981) calculated that the emplacement occurred between about 4,400 and 3,100 years BP, and that the homogenite was deposited from sediment transport induced by the Minoan tsunami. The thick and structureless homogeneous mud was later recognized in more than fifty gravity cores in a range of contrasting settings (Cita and Aloisi 2000). Mechanisms proposed for the Minoan tsunami include the entry into the sea of both pyroclastic and debris flows propagated in all directions around the island (McCoy and Heiken 2000) along with caldera collapse combined with a large tectonic earthquake (Pararas-Carayannis 1992). To simulate the wave hydrodynamically, Minoura et al. (2000) suggested a sudden volcano collapse, caldera formation, inrush of water into the caldera, and collision of water masses

Tsunamis

501

Fig. 17.9. Palaeotsunami investigation within an archaeological excavation in St George, Thalassitis, Kamari, in eastern Thera. The box indicates a distinctive sediment layer produced by a tsunami and highlights the area shown in more detail in Figure 17.10 (photo: Gerassimos Papadopoulos.)

with the caldera wall. Their simulation resulted in wave heights of >15 m in the near-field zone and of 6–11 m in northern Crete. However, the extent of the wave inundation was only several hundred metres and, although the fishing and trading economy could have been affected by the destruction of boats and harbour installations, it can be argued that a tsunami of this size would have had little long-term influence on the Minoan civilization. Another large tsunami was generated during the eruption of Columbo, a submarine volcanic edifice lying 7.3 km to the north-east of Thera (Figure 17.8). The main volcanic activity began on 26th September 1650, while volcanism on 30 September was followed by a pause in activity. During this pause a sea swell encircled the whole of Thera island and the tsunami inundated the eastern coast and swept away churches, enclosures, boats, trees, and agricultural land. On the east and west coast of Patmos island and on Ios island,

tsunami run-up heights of 30, 50, and 16 m respectively were reported. Ships and fishing boats moored at Iraklion were swept violently offshore, while vessels were crushed when the wave overtopped the city walls. The volcanic and seismic quiescence that prevailed before the tsunami struck implies that it was generated by submarine landsliding or collapse of the volcanic cone rather than by a strong earthquake or volcanic explosion (Dominey-Howes et al. 2000). A geological record of the 1650 tsunami has been recognized from the coastal site of St George’s near Kamari village in eastern Thera (Figures 17.9 and 17.10). However, Dominey-Howes et al. (2000) were unable to trace any tsunami deposit signature across three trenches, one of them being only 500 m from the St George’site. The non-volcanic clasts they found were angular to very angular in form and it is therefore thought that the non-volcanic sediments may reflect local colluvial processes. Thus, alternative hypotheses involving discontinuous sediment deposition

502

Gerassimos Papadopoulos

sides of the Amorgos basin as well as sea floor sediment instability and a geologically very recent slump (24 × 6 km in area and 3.6 × 106 m3 in volume) occupying part of the basin (Perissoratis and Papadopoulos 1999). The proximity of the slumped area to the earthquake epicentre implies seismic ground accelerations much higher than the minimum values required to initiate slumping. The slump episode may have occurred in association with the 1956 earthquake. Numerical simulations showed a discrepancy by a factor of 3–10 between the maximum reported (30 m) and simulated (3–10 m) wave amplitudes at the source region (e.g. Yalçiner et al. 1993). Field observations and interviews with eyewitnesses (Dominey-Howes 1996; Papadopoulos et al. 2005) showed that the wave amplitude may not have exceeded about 15 m in Astypalaea. Therefore, part of the discrepancy could be explained by an overestimation of the initially reported wave height. Nonetheless, a significant discrepancy remains unexplained. Thus it may be concluded that an adequate reproduction of the near-field wave amplitudes requires not only co-seismic sea-floor fault displacement but also an additional tsunamigenic component such as that from a co-seismic, massive submarine sediment slump.

Fig. 17.10. Detail of the tsunami deposits exposed in the section shown in Figure 17.9 and attributed to the 30 September 1650 event. The tsunami deposits are the dark layer behind the pen in the middle of the photograph (photo: Gerassimos Papadopoulos).

The Gulf of Corinth

and overestimation of the event’s magnitude have been considered (ibid.). The most recent large Mediterranean tsunami occurred in the Cyclades (Figure 17.3a) on 9 July 1956 after a Ms 7.4 crustal earthquake associated with normal faulting (Papadopoulos and Pavlides 1992; Chapter 16). The tsunami-generating rupture was about 100 km in length within the NE–SW trending basin formed by the Thera, Amorgos, and Astypalaea islands (Papazachos et al. 1985). Initial estimates of the near-source wave height varied between 15 and 30 m in Amorgos and Astypalaea (Galanopoulos 1957; Ambraseys 1960) (Figure 17.8). Four people drowned and extensive destruction was noted in port facilities and in both small and large vessels as well as on cultivated land and other property. Two interrelated problems arise out of the study of the 1956 tsunami. The first is the generation mechanism and the second is that numerical simulations have failed to reproduce the observed wave heights accurately. From tide gauge records Galanopoulos (1957) and Ambraseys (1960) concluded that the wave was probably produced by co-seismic landslides. Submarine geophysical survey showed normal faulting in the

The Gulf of Corinth is especially prone to tsunamis due to high seismicity, steep bathymetry and a susceptibility to coastal landsliding (Papadopoulos 2003b; Papadopoulos 2007a). It is shown on Figure 17.3a as having the highest tsunami potential in the Mediterranean region. In 373 BC it has been argued that the town of Helike, located about 7 km east of modern Aeghion (Figure 17.11a), was destroyed by an earthquake and tsunami (Guidoboni et al. 1994; Papadopoulos 1998; and see Chapter 16 for further discussion). Ten Spartan ships at anchor close by were destroyed. The lethal earthquakes of 25 May 1748 and 23 August 1817 generated similar tsunamis in Aeghion causing human losses and extensive damage to vessels, port facilities, and cultivated land. The June 1402 tsunami was also of high intensity and followed a large, possibly near-shore earthquake (Papadopoulos et al. 2000). Interestingly, the near-shore earthquake of 26 December 1861, having a magnitude comparable to the previous events, produced a tsunami of much lower intensity. Seismically triggered earth slides caused local tsunamis along the north coast on 11 June 1794, 6 July 1965, 11 February 1984, and 15 June 1995. An aseismic tsunami generated by sediment slumping at

Tsunamis

503

50 km

(a) 426 BC

M

ali

ak os

GREECE

Gu

lf

552 1817 1748,1963 Aeghion

373 BC 1402

Athens Corinth

Peloponnese

(b)

Fig. 17.11. (a) Important tsunamis in the Gulf of Corinth and the Maliakos Gulf. (b) An air photograph showing the coast to the east of the village of Psathopyrgos prior to the large landslide that produced a tsunami in the west of the Gulf of Corinth on 7 February 1963 (after Galanopoulos et al. 1964). The black line marks the extent of the sediment mass that slipped seawards. This body of sediment was approximately 1500 m in length with a maximum width of 350 m.

a river mouth hit both coasts at the western end of the Gulf of Corinth on 7 February 1963 (Figure 17.11b). The wave killed two people, injured twelve, and was responsible for serious damage to houses, cultivated land, and fishing boats (Galanopoulos et al. 1964). Numerical modelling results are consistent with run-up and inundation observations (Koutitas and Papadopoulos 1998). A similar wave of lesser intensity was observed near Aeghion on 1 January 1996.

The Maliakos Gulf The Maliakos Gulf is located on the western side of the Aegean and strong earthquake-generated tsunamis have been reported here for 426 BC and AD 551 or

552 from classical sources and archaeological evidence (Figures 17.3a and 17.11a). Records of past tsunami activity are not always consistent and care must be taken in their interpretation. For example, Papaioannou et al. (2004) suggested that the 426 BC event was, in fact, rather moderate and they argue that the large tsunami from this period may have occurred during the third century BC and that previous researchers amalgamated the two events into the earlier one at 426 BC. The Byzantine writer Procopius reported on an earthquake that struck the Gulf of Corinth in AD 551. He also described a strong tsunami in the Maliakos Gulf. However, it seems that Procopius did not actually describe a tsunami caused by the AD 552 earthquake, instead he just reproduced the classical sources detailing the 426 BC tsunami.

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Gerassimos Papadopoulos

East and North Aegean Sea A distinctive tsunami-prone region is associated with an earthquake zone around Chios Island in the eastern Aegean Sea (Figure 17.12a). Tsunamis of k 3 or 4 were observed on 20 March 1389, 12 May 1852, 8

(a) 1932

1893 Ierissos

Samothraki ~ 1300 BC

Troy

TURKEY

1856

Chios

Bosphoru s

50 km

(b)

Bosphorus

TURKEY Tekirdag

Istanbul

1343 447

1912

1999 Bay Izmit

1766

September 1852, 13 November 1856, 2 February 1866, 3 April 1881, and 23 July 1949. Based on the writings of the Greek historian Democles (fourth century BC), some catalogues list a tsunami that supposedly struck Troy in about 1300 BC. A relevant passage in the Wall War of Iliad has not attracted much attention to date. In fact, rhapsody M18–19 describes how the wall of Greeks was damaged by river flooding, yet in M26–33 it becomes clear that the sea contributed to this destruction. In fact, the wall was inundated by the sea and trees and wall materials were carried away by the water; the ground was levelled by the strong wave and the coastal zone was covered by sand; while the sea caused the flow of rivers to reverse. The vivid descriptions in the Iliad indicate that a palaeotsunami survey along the coast of Troy could potentially be of great geomorphological and archaeological interest. In the northern Aegean Sea (Figures 17.3a and 17.12a), a destructive wave that reportedly struck Potidaea on the Chalkidiki Peninsula, in April 479 BC, may imply tsunami action. This is the first tsunami described in documentary sources anywhere in the world. In AD 544 an earthquake-induced destructive inundation hit the coast of Thrace. More recently, seismic tsunamis of k 3 were observed in this area on Samothraki Island on 9 February 1893 and at Ierissos, Chalkidiki, on 26 September 1932.

Sea of Marmara

The Sea of Marmara 50 km

(c) CYPRUS

551

LEBANON

1759 1202 Akko

100 km

749

ISRAEL

JORDAN

Fig. 17.12. Important tsunamis in (a) the east and north Aegean Sea, (b) the Sea of Marmara, and (c) the Cyprus–Levantine Sea area.

This area has a long record of tsunami activity and k 3 or 4 tsunamis were caused by earthquakes on AD 26 January 447, 26 October 740, 2 September 1754, and 19 April 1878 in Izmit Bay. Tsunamis of k 3 or 4 were also caused by earthquakes on AD 25 September 478, 10 August 1265, 18 October 1343, 25 May 1419, 10 September 1509, and 10 July 1894 in Constantinople (Istanbul) as well as on 22 May 1766 in the Bosporus Straits, and on 9 August 1912 in Terkirda˘g (Figures 17.3a and 17.12b). Due to higher seismicity, the east is the most tsunami-prone side of the Marmara Sea. The large (Mw 7.4) Izmit earthquake of 17 August 1999, caused by right-lateral strike-slip faulting with a significant normal component, generated damaging waves up to 2.5 m high in De˘girmendere on the south coast, and elsewhere (Yalciner et al. 2002) with intensities up to 4 (Papadopoulos 2001). Rothaus et al. (2004) argued that the uniformity of the tsunami impact, indicating a wave coming to the south coast from 310◦ , suggests that submarine faulting was the major source of these tsunamis.

Tsunamis

Cyprus-Levantine Sea The final part of this review looks at the coasts bordering the Cyprus and Levantine Sea in the eastern Mediterranean. Archaeological excavations in Kourion, south-west Cyprus (Figures 17.3a and 17.12c), have revealed a destruction horizon attributed to the AD 21 July 365 earthquake tsunami that supposedly devastated Kourion (Soren 1988). However, an earthquake location off the shore of south-west Cyprus is not consistent with both the AD 365 seismic damage distribution and tsunami wave propagation to the south, west, and north-west of Crete (Figure 17.6). An alternative explanation has been put forward that Kourion was hit by a non-tsunami-generating earthquake around AD 370 (Ambraseys 1965; Guidoboni et al. 1994). In Cyprus, earthquake tsunamis were reported on 22 May 1201 or 1202, 11 May 1222, and 10 September 1953, the first two being of high intensity. Geomorphic evidence along coastal sections of southern Cyprus and radiocarbon dating results indicate tsunami activity between 1530 and 1821 (Whelan and Kelletat 2002). In the left-lateral strike-slip Levantine rift, tsunamigenerating earthquakes have been identified from historical records (Figures 17.3a and 17.12c). On 9 July 551 the sea retreated for a mile and many ships were destroyed along the coasts of Lebanon, Syria, and Palestine, while after an earthquake on AD 18 January 749, the waves ‘rose up to the sky’ and destroyed most of the cities and villages along the coasts of Israel, Palestine, and Syria (Guidoboni et al. 1994, Darawcheh et al. 2000). Similar tsunamis were reported after strong earthquakes on 5 December 1033 (that shook the region around the Jordan Valley), on 29 May 1068 (an event possibly centred to the south of the Levantine rift), and on 14 January 1546 near the Dead Sea (Ambraseys et al. 1994).

Tsunami Generation Mechanisms The large propagation areas of waves such as those of AD 365 and 1303 events leave little doubt that they were produced by co-seismic, sea-bottom dip-slip faulting. However, with only limited data about other tsunamis associated with large earthquakes of dip-slip faulting—such as the 1908 event in the Messina straits and the 1956 event in the Cyclades—it is often difficult to distinguish between co-seismic, landslide, or combined source mechanisms. In addition, the generation of tsunamis with strong local effects by earthquakes occurring on land—such as along the strike-slip Levantine rift—remains unexplained. Submarine landsliding

505

triggered by powerful seismic shaking is a possibility, but an alternative mechanism is the dynamic excitation of tsunamis due to seismic energy transmission to the continental shelf by the vertical component of the longperiod Rayleigh waves. However, in the highly seismogenic Cephalonia–Lefkada system of strike-slip faulting in the Ionian Sea, the tsunami activity is low—and in the Sea of Marmara, the tsunami potential is only intermediate (Figure 17.3a). This may be due to the fact that the earthquake focal mechanism involves mainly a strike-slip rather than a dip-slip component, while earth slumping may contribute to more local tsunami generation. Volcanic eruptions are much less frequent than earthquakes in the Mediterranean and only a few tsunamis are therefore attributed to volcanic activity (Chapter 15). However, two of the largest known tsunamis were generated by strong eruptions in the Thera volcanic complex—namely the seventeenth-century BC Minoan tsunami and the AD 1650 Columbo tsunami. The generation of the Minoan tsunami may have included several mechanisms including caldera collapse, strong earthquake activity, pyroclastic surges and flows, and lahars and debris flows into the sea. Additional mechanisms that are worthy of more research include tsunami generation from atmospheric pressure changes due to the volcanic eruption, as well as abrupt sea-bottom impulse following the caldera collapse. The Columbo tsunami was probably caused by the partial collapse of the submarine caldera. Local but strong tsunamis caused by landsliding during eruptive activity of Stromboli have been reported repeatedly (Chapters 1 and 15). Aseismic submarine or coastal landsliding is an important agent of locally powerful tsunamis (Papadopoulos et al. 2007a). The physiographic features of the Gulf of Corinth favour these processes and the most recent events were observed in 1963 and 1996. In the western Mediterranean in 1979, the collapse of a mass of unconsolidated artificial embankment into the sea in Nice on the Côte d’Azur, caused a local tsunami of similar character to that of 1963 in the Gulf of Corinth. Both events produced a large amplitude wave leading to loss of life and significant destruction in the near-source coastal zone.

Tsunami Hazard This section evaluates tsunami hazard in Mediterranean coastal regions using four independent approaches. The first is a comparative description of the spatial distribution of tsunami events based on the geological

Tsunami frequency (n)

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30

70

25

60 50

20 40 15 30 10

20

5

10 0

0 3

W

0

5

10

15

20

Longitude (°)

25

30

32

35

35

40

45

47

Latitude (°N)

E

Fig. 17.13. The frequency (n) of tsunamis in the Mediterranean Sea as a function of longitude (left) and latitude (right) for the period 1628 BC to AD 2003.

and historical tsunami record over the last 3.5 millennia. Second, tsunami recurrence is calculated from empirical intensity–frequency relationships. The third is a statistical approach to estimate tsunami potential in a uniform way across the several tsunamigenic zones of the Mediterranean Sea. Finally, the occurrence of tsunamis is examined as a function of earthquake size. The data sets used here come from the various tsunami catalogues listed in the earlier sections. The known tsunami sources in the Mediterranean are concentrated in eighteen zones of high seismicity and/or volcanism as shown in Figure 17.3a (Chapters 15 and 16). In the Mediterranean as a whole, the tsunami hazard (in terms of the event count) increases gradually from west to east—again tracking the basin-wide pattern of seismicity—although a decrease is apparent in the easternmost part of the basin. In fact, maximum activity occurs between 20◦ E and 30◦ E in the complex tectonic terrain and highly seismically-active structures of Greece (Figure 17.13) including the western and eastern segments of the Hellenic arc and the Gulf of Corinth. Significant tsunami activity is also concentrated in the seismic and volcanic region of Italy between 15◦ E and 17◦ E. The main belt of tsunami activity lies within a latitudinal zone between approximately 37◦ N and 41◦ N, which includes Greece, the Marmara Sea, the Tyrrhenian Sea, and southern Italy (Figure 17.13). To determine the frequency of tsunami occurrence, some authors have used the intensity–frequency relationship: log Nc = ac − bk

(1)

where Nc is the cumulative number of events of intensity equal to or larger than k observed in a time interval of c years, and a c and b are parameters determined by the data. Equation (1) is analogous to the classic earthquake magnitude–frequency relationship (Gutenberg and Richter 1944) and applies to tsunamiprone regions where relatively good datasets are available. For c = 1 formula (1) becomes: log N = a − bk

(2)

a = a c − log c

(3)

where:

Then, the mean tsunami recurrence for intensity ≥k is equal to: T = 10b k−a

(4)

The most likely maximum tsunami intensity (kt ) expected in a time interval of t years is given by: kt = (a + log t)/b

(5)

However, the small number of data points does not allow the application of this relationship in most of the tsunami-prone areas of the Mediterranean. Thus, this approach may only be applied in the larger tsunamigenic regions such as Greece, Italy, and the entire Mediterranean Sea. From the temporal distribution of the reported tsunami events a complete record for events of kc ≥ 3 from AD 1600 onwards is assumed (Figure 17.14). The results are shown in Figure 17.15 and Table 17.2 and suggest that, in the entire Mediterranean, the mean recurrence interval for tsunamis of

Tsunamis 300

2.5

(a)

(a) Greece

250

2.0

n =13.869e

200

log frequency (n)

Tsunami frequency (n)

507

0.001t

2

R= 0.938

150 100

1.5 log n = 3.67–0.61k

1.0

r2 = 0.996 0.5

50

0 0

0 2000

1000

1000

0

1

2

3

4

5

6

2.5

3000

2000

(b) Italy

Years (BC/AD)

log frequency (n)

2.0 6

Tsunami intensity (k)

(b) 5 4

1.5 1.0 log n = 2.16–0.31k

0.5

3

r2 = 0.972 0

2

0

1

2

3

4

5

6

2.5 1000

500

500

0

1000

(c) Mediterranean

2000

1500

Years (BC/AD)

Fig. 17.14. (a) The cumulative frequency (n) of Mediterranean Sea tsunamis as a function of time for the period 1628 BC to AD 2003. (b) The intensity (k) of tsunamis between 1500 BC and AD 2000.

2.0

log frequency (n)

1 1500

1.5 log n = 3.56–0.51k

1.0

r2 = 0.999 0.5

kc ≥ 3, 4, 5, and 6 is 4, 12, 40, and 130 years, respectively. This analysis also shows that Greece is characterized by a higher tsunami frequency than Italy, with the exception of k 6 events, which appear to recur more frequently in the Italian region. This last outcome, however, should be treated with some caution since it is based only on the post-1600 statistics and some very significant high-intensity events, like the AD 365 and 1303 tsunamis in the Hellenic arc and the 1402 event in the Gulf of Corinth, were not considered in the calculation. TABLE 17.2. Mean return period (T ) of tsunami intensity (k) and the most likely maximum tsunami intensity (kt ) to be observed in time interval t for various parts of the Mediterranean Sea Region

Mediterranean Sea Greece Italy Gulf of Corinth*

T (years)

kt

k≥3

4

5

6

t (yrs) = 1

10

100

4 6 26 40

13 24 55 103

41 98 115 261

132 399 242 662

2 2

4 4 2 2

6 5 5 4

* Modified from Papadopoulos (2003b).

0 0

1

2

3

4

5

6

Tsunami intensity (k)

Fig. 17.15. Diagrams showing the relationship between intensity and frequency of tsunamis from the database for (a) Greece, (b) Italy, and (c) the Mediterranean Sea.

Since detailed intensity-frequency relationships are impossible to obtain in most tsunami-prone regions due to small datasets, an alternative procedure was introduced for the uniform assessment of tsunami potential. This procedure combines information on the frequency and intensity of events, but does not consider the potential hazard to far-field locations which is due to wave propagation in remote places. Thus, the tsunami potential or hazard (H ) of a particular zone or area is defined as the normalized quantity: H = Ha /Hmi n

(6)

where the absolute potential, Ha , is an increasing function of the weighted event intensity and the event

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frequency: Ha =

n 

(kci × jc )/tc

i =1

where Hmi n = min {Ha }, kci is the intensity of event i, and jc is a weighted factor of kc . Factor jc has been defined to follow a power law of base 2:21 for kc = 3, 22 for kc = 4, 23 for kc = 5, and 24 for kc = 6. Only events of kc ≥ 3 were considered. The lower date of the time interval (tc ) over which the data set is complete is variable, while the upper date was fixed at the end of 2003. The earliest event in the record was assumed to be AD 300 for kc = 6, AD 1000 for kc = 5, AD 1600 for kc = 4, and AD 1750 for kc = 3 (Figure 17.14). Equation (6) produces a relative tsunami potential scale with a minimum value equal to 1. The eighteen tsunamigenic zones in the Mediterranean were classified into four groups of potential and these are shown in Figure 17.3a and Table 17.3. The lowest potential was found in Tuscany (H = 1), while low potential characterizes the coastal region from Liguria to the Côte d’Azur and the north Aegean Sea. The potential may be described as intermediate in the Alboran Sea, the Aeolian Islands, the Tyrrhenian Calabria, eastern Sicily, and the Messina Straits, the Gargano promontory, the coasts of Albania and Montenegro, the Cyclades, the eastern Aegean Sea, the Sea of Marmara, and the Levantine Sea. High potential

TABLE 17.3. Tsunami potential in each of the tsunamigenic zones of the Mediterranean shown in Figure 17.3a No in Fig. 17.3a 3 7 2 15 6 11 8 16 5 18 1 4 14 9 10 12

Region Tuscany Gargano, Italy Liguria and the Côte d’Azur North Aegean Sea East Sicily and Messina Straits Cyclades Islands Eastern Adriatic Sea Sea of Marmara Tyrrhenian Calabria Levantine Sea Alboran Sea Aeolian Islands Eastern Aegean Sea Western Hellenic arc Eastern Hellenic arc Gulf of Corinth

Ha

H

0.024 1 0.040 1.67 0.047 1.96 0.047 1.96 0.096 4.0 0.113 4.71 0.119 4.96 0.126 5.25 0.127 5.29 0.127 5.29 0.134 5.58 0.134 5.58 0.142 5.92 0.199 8.29 0.223 9.29 0.308 12.83

Potential low low low low intermediate intermediate intermediate intermediate intermediate intermediate intermediate intermediate intermediate high high very high

Note: Ha = absolute potential, H = normalized potential. Very low potential has been tentatively assigned to Campania, Ionian Apulia and Calabria, the central and northern Ionian Sea, the Maliakos Gulf, and Cyprus. See text for further discussion.

values were determined for the western and eastern segments of the Hellenic arc, while the Gulf of Corinth is classified as having the highest tsunami potential (H = 12.83) in the Mediterranean Sea (Figure 17.3a; Chapter 13). In such regions as Cyprus, the Maliakos Gulf, the central and northern Ionian Sea, as well as Ionian Apulia and Calabria, only a few strong or moderate tsunamis have been documented—but these took place before the time period of complete data sets used in the above analysis. A characteristic example is the tsunami wave of intensity k 4 of AD 1222 in Cyprus (Fokaefs and Papadopoulos 2007). In these regions the tsunami potential is not negligible, but it is impossible to quantify at present using the methods outlined above. Thus, these regions are provisionally classified as being of very low tsunami potential. The tsunami database for much of the North African coast of Tunisia, Libya, and Egypt is also very patchy and this probably reflects a combination of low tsunami potential (due to lower relief and low seismicity) and limited documentary and archaeological records. It is important to appreciate that the intensity of an individual tsunami depends on many factors including earthquake size and epicentral location as well as the nature of the crustal displacement (Lorito et al. 2008) The local and regional sea floor bathymetry is also important—as is the geomorphology of the coastal zone receiving the tsunami wave. Given the range of variables, it is not surprising that there is not a strong correlation between tsunami intensity and earthquake magnitude and intensity (Figure 17.16). However, it is possible to identify an upper bounding envelope in both diagrams which shows that tsunami intensity does not exceed a certain value unless it is generated by an earthquake that exceeds a minimum size. For example, to produce a tsunami with an intensity ≥ 4 requires at least an earthquake magnitude of 6.3 or an earthquake intensity of 7.

Risk Mitigation Technology Tsunami risk mitigation can be achieved through a series of actions including: 1. the operation of instrumental early warning systems; 2. the construction of breakwaters; 3. the development of risk mapping using GIS tools; 4. the implementation of dedicated civil protection plans and public education.

Tsunamis

(a)

(b) 6

6

0.41M

0.36i

k = 0.337e r2 = 0.999

k = 0.313e

5

2

Tsunami intensity (k)

5

Tsunami intensity (k)

509

4 3 2 1

r = 0.996 4 3 2 1 0

0

IV

V

VI

VII

VIII

IX

X

XI

XII

Earthquake intensity (i )

5

6

7

8

Earthquake magnitude (M)

Fig. 17.16. Tsunami intensity (k) as a function of (a) earthquake intensity and (b) earthquake magnitude for the entire Mediterranean Sea database. Upper bounding envelopes are shown.

Although the countries bordering the Mediterranean Sea do not have an established tradition in this field (in contrast to Japan, for example, and other circumPacific countries), significant progress has been made over the last fifteen years or so. For example, experimental tsunami warning systems, consisting of both seismic and tide-gauge instruments, have recently been tested in the Messina Straits and the southern Aegean Sea (Piscini et al. 1998, Papadopoulos 2003c). After the strong tsunami of 30 December 2002 in Stromboli, the civil protection authorities established a local tsunami alarm system. A major problem, however, for tsunami risk mitigation in the Mediterranean Sea is that travel times from tsunami sources to threatened coastal communities are short, ranging from a few minutes to around one hour depending on the source location. However, operationally viable systems can be achieved by minimizing the alarm time to no more than about five minutes from the earthquake origin time. In 2005 the Intergovernmental Oceanographic Commission of UNESCO initiated a project of establish an instrumental early warning system for tsunamis in the Mediterranean Sea and the northeast Atlantic Ocean. Pilot studies of risk mapping on a microzonation basis appear to be useful for long-term planning in the mitigation of tsunami risk (Papadopoulos and Dermentzopoulos 1998; Papathoma et al. 2003). In Japan, the construction of breakwaters has been used for the protection of some coastal zones from tsunami attack, but such measures have never been recommended in the Mediterranean Sea. In comparison to other natural hazards in the Mediterranean such as earthquakes,

volcanoes, and large fluvial floods, there has only been limited progress to date in the development of civil protection plans and raising public awareness of the potential threat from tsunamis in the region. While public awareness of the tsunami hazard more generally has undoubtedly been raised following the catastrophic tsunami in the Indian Ocean in December 2004, awareness of the extent of the local tsunami threat in the Mediterranean region is still rather limited. Progress in the types of action outlined above is very important for the development of an effective tsunami risk mitigation policy for the Mediterranean Sea coasts. Using tsunami-wave simulations generated from mathematical models, Lorito et al. (2008) have calculated a range of parameters including wave height profiles and wave travel times to assess the potential threats to southern Italy from tsunamis generated by earthquakes in three different source zones, namely the southern Tyrrhenian thrust belt, the Tell-Atlas thrust belt, and the western Hellenic arc. The outcomes show a highly variable impact for tsunamis produced in the different source zones and this analysis highlights the need for scenario testing to be undertaken at the scale of the entire Mediterranean basin (Lorito et al. 2008).

Concluding Remarks Segments of the Mediterranean Sea coast have been struck in the past by large, destructive tsunamis generated by submarine earthquakes and volcanic eruptions (e.g. Soloviev et al. 2000), while powerful tsunamis

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have also been locally generated by landslides (e.g. Papadopoulos 1993b). The recurrence interval for tsunami intensities ≥ 4, 5, and 6 is of the order of 12, 40, and 130 years respectively. In this chapter, the highest tsunami potential has been calculated in the Hellenic arc and the Gulf of Corinth. However, infrequent but large events have been recognized for the Cyclades in the southern Aegean Sea and the Straits of Messina in southern Italy, as well as in the Alboran, Levantine, and Marmara Seas. It is clear, therefore, that tsunami waves should not be neglected as a potential source of risk that threaten coastal communities of the Mediterranean. Several approaches offer considerable potential to improve our understanding of the tsunami phenomena and associated hazards including the identification of earthquake, volcanic eruption, and landslide sources and mechanisms as well as palaeotsunami investigations. To improve tsunami risk assessment, more work is needed to extend the event database, to simulate wave generation, propagation, and coastal inundation using numerical models and to map the components of physical and anthropogenic risk with the use of GIS tools. Recent events in south-east Asia and the Indian Ocean margins have highlighted the importance of developing risk mitigation strategies that include public awareness activities, the development of instrumental tsunami warning systems and the elaboration of specific civil protection plans across the Mediterranean region.

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Tsunamis in the Western Corinthos Bay. Proceedings of the International Conference on Tsunamis, Paris, 26–8 May, 247–54. Lorito, S., Tiberti, M. M., Basili, R., Piatanesi, A., and Valensise. G. (2008), Earthquake-generated tsunamis in the Mediterranean Sea: Scenarios of potential threats to Southern Italy. Journal of Geophysical Research 113, B01301, doi: 10.1029/2007JB004943. McCoy, F. W. and Heiken, G. (2000), Tsunami generated by the Late Bronze Age eruption of Thera (Santorini), Greece. Pure and Applied Geophysics 157: 1227–56. Marinatos, S. (1939), The volcanic destruction of Minoan Crete. Antiquity 13: 425–39. Marinos, G. and Melidonis, N. (1971), On the strength of seaquakes (tsunamis) during the prehistoric eruptions in Santorini, in Proceedings of the 1st International Scientific Congress on the Volcano of Thera, Athens, 15–23 September 1969, 277–88. Minoura, K., Imamura, F., Kuran, U., Nakamura, T., Papadopoulos, G. A., Takahashi, T., and Yalçiner, A. C. (2000), Discovery of Minoan tsunami deposits. Geology 28: 59–62. Monserrat, S., Ibbetson, A., and Thorpe, A. J. (1991), Atmospheric gravity waves and the ‘Rissage’ phenomenon, Quarterly Journal of the Royal Meteorological Society 117: 553–70. Murty, T. S. (1977), Seismic Sea Waves–Tsunamis. Bulletin of the Fisheries Research Board Canada 198: 1–337. and Loomis, H. G. (1980), A new objective tsunami magnitude scale. Marine Geodesy 4: 267–82. Neri, G., Barberi, G., Oliva, G., and Orecchio, B. (2004), Tectonic stress and seismogenic faulting in the area of the 1908 Messina earthquake, south Italy. Geophysical Research Letters 31: 1–5. Papadopoulos, G. A. (1993a), On some exceptional seismic (?) sea-waves in the Greek Archipelago. Science of Tsunami Hazard, 11, 25–3. (1993b), Seismic faulting and non-seismic tsunami generation in Greece. Proceedings of the IUGG/IOC International Tsunami Symposium, Wakayama, Japan 23–7 August, 1993, 115–22. (1998), A reconstruction of the 373 B.C. large earthquake in the western Corinthos Gulf. Proceedings of the 2nd International Conference on Ancient Eliki, Aeghion, Athens, 1–3 Dec. 1995, 479–94. (2001), Tsunamis in the East Mediterranean: A catalogue for the area of Greece and adjacent seas. Proceedings of the Workshop on Tsunami Risk Assessment Beyond 2000: Theory, Practice, Plans. Moscow, 14–16 June 2000, 34–42. (2003a), Quantification of tsunamis: a review, in A. C. Yalçiner et al. (eds.), Submarine Landslides and Tsunamis. Kluwer, Dordrecht, 285–91. (2003b), Tsunami hazard in the Eastern Mediterranean: Strong earthquakes and tsunamis in the Corinth Gulf, Central Greece. Natural Hazards 29: 437–64. (2003c), A tsunami warning system in the SW Aegean Sea, Greece, in J. Zschau and A. N. Küppers (eds.), Early Warning Systems for Natural Disaster Reduction. Springer, New York, 549–52. and Chalkis, B. G. (1984), Tsunamis observed in Greece and the surrounding area from antiquity up to present times. Marine Geology 56: 309–17. and Dermentzopoulos, Th. (1998), A tsunami risk management pilot study in Heraklion, Crete Isl. Greece. Natural Hazards 18: 91–118.

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and Imamura, F. (2001), A proposal for a new tsunami intensity scale. Proceedings of the International Tsunami Symposium 2001, Seattle, 7–10 August, 569–77. and Kortekaas, S. (2003), Characterisics of landslide gererated tsunamis from observational data, in L. Locat and J. Mienert (eds.), Submarine Mass Movements and their Consequences. Kluwer, Dordrecht, 367–74. and Pavlides, S. B. (1992), The large 1956 earthquake in the South Aegean: Macroseismic field configuration, faulting, and neotectonics of Amorgos island. Earth and Planetary Science Letters 113: 383–96. and Vassilopoulou, A. (2001), Historical and archaeological evidence of earthquakes and tsunamis felt in the Kythira strait, Greece, in G. T. Hebenstreit (ed.), Tsunami Research at the End of a Critical Decade. Kluwer, Dordrecht, 119–38. and Plessa, A. (2000), A new catalogue of historical earthquakes and tsunamis in the Corinth rift, Central Greece: 480 B.C.–A.D. 1910, in G. A. Papadopoulos (ed.), Historical Earthquakes and Tsunamis in the Corinth Rift, Central Greece. Publication 12, Institute of Geodynamics, National Observatory of Athens, 9–119. Imamura, I., Minoura, K., Takahashi, T., Kuran, U., and Yalciner, A. (2004), Strong earthquakes and tsunamis in the East Hellenic Arc. Geophysical Research Abstracts 6: 3212. Karakatsanis, S., Fokaefs, A., Orfanogiannaki, K., Daskalaki, E., and Diakogianni, G. (2005), The 9 July 1956 large tsunami in the south Aegean Sea: compilation of a data basis and re-evaluation, in G. A. Papadopoulos and K. Satake (eds.), Proceedings of the 22nd International Tsunami Symposium, Chania, 27–9 June, 173–80. Papadopoulos, G. A., Daskalaki, E., and Fokaefs, A. (2007a), Tsunamis generated by coastal and submarine landslides in the Mediterranean Sea, in V. Lykousis, D. Sakellariou, and J. Locat (eds.), Submarine Mass Movements and their Consequences. Advances in Natural and Technological Hazards Research 27. Springer, Dordrecht, 415–22. and Giraleas, N. (2007b), Tsunami hazards in the Eastern Mediterranean: strong earthquakes and tsunamis in the East Hellenic Arc and Trench system. Natural Hazards and Earth System Sciences 7: 57–64. Papaioannou, I., Papadopoulos, G. A., and Pavlides, S. (2004), The earthquake of 426B.C. in N. Evoikos Gulf revisited: amalgamation of two different strong earthquake events? Bulletin of the Geological Society of Greece 36: 1477–81. Papathoma, M., Dominey-Howes, D., Zong, Y., and Smith, D. (2003), Assessing tsunami vulnerability, an example from Herakleio, Crete. Natural Hazards and Earth System Sciences 3: 1–13. Papazachos, B. C., Koutitas, C. H., Hatzidimitriou, M. P., Karacostas, G. B., and Papaioannou, A. Ch. (1985), Source and short-distance propagation of the July 9, 1956 southern Aegean tsunami. Marine Geology 65: 343–51. Pararas-Carayannis, G. (1992), The tsunami generated from the eruption of the volcano of Santorini in the Bronge Age. Natural Hazards 5: 115–23. Patacca, E. and Scandone, P. (2004), The 1627 Gargano earthquake (Southern Italy): Identification and characterization of the causative fault. Journal of Seismology 8: 259–73. Perissoratis, C. and Papadopoulos, G. A. (1999), Sediment instability and slumping in the southern Aegean Sea and the case history of the 1956 tsunami. Marine Geology 161: 287–305.

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Pirazzoli, P. A., Ausseil-Badie, J., Giresse, P., Hadjidaki, E., and Arnold, M. (1992), Historical environmental changes at Phalasarna harbour, west Crete. Geoarchaeology 7: 371–92. Piscini, A., Maramai, A., and Tinti, S. (1998), Pilot local tsunami warning system in Augusta, Eastern Sicily, Italy. International Conferernce on Tsunamis, Paris, May 26–8 1998, 137–48. Rothaus, R. M., Reinhardt, E., and Noller, J. (2004), Regional considerations of coastline change, tsunami damage and recovery along the southern coast of the bay of Izmit—the Kocaeli (Turkey) earthquake of 17 August 1999. Natural Hazards 31: 233–52. Shaw, B., Ambraseys, N. N., England, P. C., Floyd, M. A., Gorman, G. J., Higham, T. F. G., Jackson, J. A., Nocquet, J. M., Pain, C. C., and Piggott, M. D. (2008), Eastern Mediterranean tectonics and tsunami hazard inferred from the AD365 earthquake. Nature Geoscience 1: 268–76. Shuto, N. (1993), Tsunami intensity and disasters, in S. Tinti (ed.), Tsunamis in the World. Kluwer, Dordrecht, 197–216. Sigurdsson, H., Carey, S., and Devine, J. D. (1990), Assessment of mass, dynamics and environmental effects of the Minoan eruption of Santorini volcano, in: D. A. Hardy et al. (eds.), Thera and the Aegean World III. Thera Foundation, London, ii. 100–12. Soloviev, S. L. (1970), Recurrence of tsunamis in the Pacific, in W. M. Adams (ed.), Tsunamis in the Pacific Ocean. East–West Center Press, Honolulu, 149–63. Solovieva, O., Go, C., Kim, K., and Shchetnikov, A. (2000), Tsunamis in the Mediterranean Sea 2000 B.C. to 2000 A.D. Kluwer, Dordrecht. Soren, D. (1988), The day the world ended at Kourion— Reconstructing an ancient earthquake. National Geographic, July, 30–53. Thommeret, Y., Thommeret, J., Laborel, J., Montaggioni, L. F., and Pirazzoli, P. A. (1981), Late Holocene shoreline changes

and seismo-tectonic displacements in western Crete (Greece). Zeitschrift für Geomorphologie, Suppl. 40: 127–49. Tinti, S. and Guidoboni, E. (1988), Revision of the tsunamis occurred in 1783 in Calabria and Sicily (Italy). Science of Tsunami Hazards 6: 17–22. and Maramai, A. (1996), Catalogue of tsunamis generated in Italy and in Côte d’Azur, France: a step towards a unified catalogue of tsunamis in Europe. Annali di Geofisica 39: 1253–99. Maramai, A., and Favali, P. (1995), The Gargano promontory: An important Italian seismogenic-tsunamigenic area. Marine Geology 122: 227–41. (2001), Contribution of tsunami data analysis to constrain the seismic source: the case of the 1693 eastern Sicily earthquake. Journal of Seismology 5: 41–61. and Graziani, L. (2004), The new catalogue of Italian tsunamis. Natural Hazards 33: 439–65. Valensise, G. and Pantosti, D. (1992), A 125 kyr-long geological record of seismic source repeatability: the Messina Straits (southern Italy) and the 1908 earthquake (Ms 71/2). Terra Nova 4: 472–83. Van Dorn, W. G. (1968), Tsunamis. Contemporary Physics 9: 145–64. Whelan, F. and Kelletat, D. (2002), Geomorphic evidence and relative and absolute dating results for tsunami events on Cyprus. Science Tsunami Hazards 20: 3–18. Yalçiner, A. C., Kuran, U., Akyarli, A., and Imamura, F. (1993), An investigation on the propagation of tsunamis in the Aegean sea by mathematical modeling, in Proceedings of the IUGG/IOC International Tsunami Symposium, Wakayama, 23–7 August, 65–75. Alpar, B., Altinok, Y., Özbay, I., and Imamura, F. (2002), Tsunamis in the Sea of Marmara—Historical documents for the past, models for the future. Marine Geology 190: 445–63.

This chapter should be cited as follows Papadopoulos, G. A. (2009), Tsunamis, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 493–512.

18

Storms and Floods María del Carmen Llasat

Introduction Floods are the most common natural hazard in the Mediterranean region and they result in the greatest economic losses. For regions such as eastern Spain, the south of France, Italy, and the west of the Balkan Peninsula, it can be argued that damaging floods are not infrequent and they can be considered a key component of the climatic and hydrological regime (Barnolas and Llasat 2007). The economic and social impacts of flooding are not just a function of the high frequency of large floods, but also of the vulnerability created by various human activities. In common with other parts of the Mediterranean, the regions listed above are characterized by widespread and intensive economic activity as well as high population densities—often in river valleys and on coastal plains—and this combination results in very significant losses following major flood events (Ruin et al. 2008). Floods and severe weather events affect all parts of the Mediterranean, but their frequency and impact are not homogeneous across the region. Between 1991 and 1995 the damage caused by floods in Mediterranean Europe amounted to €80 billion (Estrela et al. 2000). Table 18.1 shows the main flood events that have occurred in the European Mediterranean since 1990. Although high-magnitude wind storms are not as important as heavy rains and floods, they are considered to be the second most important meteorological hazard in the Mediterranean region because of their frequency and the damage caused both inland and to sea traffic. In addition to the wind storms produced by well-known winds such as the bora or sirocco, severe-weather events and storms produced by deep lows can affect the northern and southern parts of the Mediterranean basin. The most catastrophic wind storm in recent years was the November 2001 event that caused over 600 deaths in

Algeria and four deaths—and damage put at over €37 million—in the Balearic Islands of Spain. The flooding and wind storms that occur in the Mediterranean and adjacent countries are intrinsically related to both Mediterranean meteorology itself and the marked cyclogenesis recorded in the region (Chapter 3). This chapter begins with a presentation of the forcing factors and, in particular, the role that the region’s physical geography plays in generating such adverse weather phenomena. The section devoted to heavy rainfall events takes account of recent advances in the classification of the convective systems associated with them, based on meteorological radar information (Steiner et al. 1995; Rigo and Llasat 2004; Barnolas et al. 2008). Certain climatic features are included as well as the main meteorological configurations that lead to heavy rainfall episodes, with reference to their synoptic, thermodynamic, and mesoscalar aspects. Detailed discussion of fluvial floods begins with a classification of the types of flooding found in the Mediterranean zone and a presentation of the climatology in relation to both the spatial and temporal distribution of floods. The climatic aspects are linked to a discussion of the long-term record of flooding in the Mediterranean region based on research into historical climatology. Following a section devoted to wind storms, the chapter closes with a short section on tornadoes and hailstorm events in Mediterranean countries.

Cyclones and Meteorological Hazards in the Mediterranean Region It can be argued that the Mediterranean region is geographically and meteorologically one of the most clearly defined regions in the world, with a physical framework that creates a specific Mediterranean air mass and a distinctive Mediterranean meteorology (Jansà 1966,

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María del Carmen Llasat TABLE 18.1. Major flood events in the European Mediterranean since 1990 Date 1 October 1990 15 November 1990 2 June 1992 22–3 September 1992 26–8 September 1992 3–6 October 1992 31 October 1992 November 1992 23 September 1993 1 November 1993 January 1994 October 1994 10 October 1994 4–6 November 1994 1994 19–25 April 1995 11 August 1995 19 September 1995 4–6 October 1995 19 June 1996 8 July 1996 7 July 1996 7–10 October 1996 8 October 1996 14 October 1996 1997 1998 3–5 May 1999 12–13 November 1999 10–14 June 2000 September 2000 13–16 October 2000 19–25 October 2000 6 and 23 November 2000 9–15 November 2001 8–9 September 2002 29–1 December 2002

Location Southern France Crotone, Italy South-east France Vaisson-la-Romaine, France; Savona, Italy Languedoc-Rousillon, France; Genoa, Italy Veneto, Italy Costa del Tirreno, Sicily Crotone, Italy Liguria, Italy Corcega, France Francia Athens, Greece Catalonia, Spain Piemonte, Italy Greece Piemonte, Italy La Ciotat, France Friuli, France Nîmes, France; Liguria, Italy Versilia, Italy Piemonte, Italy Biescas, Spain Piemonte, Italy Emilia-Romagna, Calabria, Italy Crotone, Italy Badajoz, Spain Sarno, Italy Piemonte, Italy Aude, France; Sardinia, Italy; Catalonia, Spain Catalonia, Spain; Piemonte, Italy Soverato, Italy Piemonte, Val d’Aosta, Liguria, Italy Catalonia, Murcia, Spain Liguria, Italy Algeria Balearic Islands, Spain Gard and other departments, France Lombardia, Italy

1997). The region is surrounded by an almost continuous barrier of mountains that hinder air mass movements. Its position between the cold Euro-Asiatic lands and the warm lands of Africa allows extra-tropical and subtropical influences to alternate and interact (Chapter 3). In addition, the warm surface and deepest water of the Mediterranean Sea favours the interchange of latent and sensible heat (Chapter 2). As a result there are two primary meteorological effects that arise almost continuously (Jansà 1997). First, the formation of a particular low-level Mediterranean air mass, around 1,500–2,000 m in thickness, that is warm and wet from autumn to spring and relatively cold and wet

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in summer, with the coastlines and the mountain barriers forming almost permanent frontal boundaries. As a result its vertical thermodynamic profile commonly shows potential convective instability. The cooling of the upper layer and/or the lifting of the whole air column can release this latent convective energy and give rise— if the other conditions are present—to heavy rainfall (Llasat and Puigcerver, 1992). Second, the emergence of lee depressions is important as well as the creation of low-level potential vorticity nuclei because of the interaction of the airflow with the mountainous barriers. These can affect the local winds and the production and distribution of rainfall.

Storms and Floods

It is helpful at this point to summarize briefly these meteorological effects from a natural hazards perspective. The Mediterranean region sees the highest concentration of cyclogenesis in the world, the heaviest extra-tropical rainfall events (as much as 800 mm in 24 hours), and very strong local winds with sustained speeds of 20 or even 25 m s−1 (Jansà 1997). In view of these factors, the World Meteorological Organization (WMO) has promoted various programmes on Mediterranean cyclogenesis, such as the Mediterranean Cyclones and Adverse Weather Phenomena Study Project (WMO 1995) or MEDEX—Mediterranean Experiment (Jansà 2002). In most cases, the adverse weather phenomena are connected with both the Mediterranean Sea and the surrounding land and mountain ranges. A good deal of the research into severe weather meteorology and related phenomena in the region has been focused on the areas shown in Figure 18.1a. The area shown in Figure 18.1b will be used to illustrate some examples of high magnitude storms and floods at various points in this chapter. The recent climatology of surface cyclones covering the whole Mediterranean region has been investigated from an operational analysis spanning June 1998 to May 2001 (Gil et al. 2003) for the area lying between 25◦ N and 49◦ N and 12◦ W and 36◦ E (Figure 18.2). This shows that the mean cyclone frequency for the West Zone is about 2,910 cyclones per year, while in the East Zone this figure is 2,248 cyclones per year. If a size filter of 200 km is used in order to eliminate the minor cyclonic perturbations, these figures are 437 and 353 respectively. Figure 18.2 shows the maximum concentrations in the Gulf of Genoa and close to Cyprus. Among other factors, these two maxima are related to the lee cyclogenesis effect produced by the Alps and the Anatolian Peninsula and the coastal mountains of Turkey. Secondary maxima are located south of the Pyrenees, in the Aegean Sea and along the Algerian coast (Figure 18.2). This geographical distribution changes over the course of the year; with summer being the season of greatest frequency and with the maximum concentration of cyclone centres in the south of the Iberian Peninsula and near Cyprus. During the summer, centres of cyclogenesis are mainly observed over land, while in winter most cyclones are located over the sea. The importance of orography in cyclones and cyclogenesis location in Mediterranean areas has been shown by Genovés et al. (1997) based on the application of the HIRLAM (High Resolution Limited Area Model) mesoscale analysis model (0.5◦ network). An initial quantitative statistical assessment of the weight of the orographic factor in Mediterranean cyclone generation shows absolute maxima in the lee of high mountain

515

ranges, where the most intense mesolows are detected. The majority of these are stationary and only a few develop into intense cyclogenesis when a high-level disturbance approaches. A particular case of atmospheric depressions that could be related to some flood events are those associated with cold air pools. According to the first definition by Scherhag (1948), a cold air pool is a deep, cold-core low that is undetectable at ground level. Preferred areas for cold pool genesis are generally associated with either high ground or warm seas (Llasat 1991). These areas shift from the eastern Atlantic in spring to a poorly defined distribution in summer, whereas in autumn and winter the preferred areas shift southwards to the Mediterranean Sea. In this region, preferred areas are to the north-east of the Balearic Islands, the northern Tyrrhenian and Adriatic seas, the centre of the Balkan Peninsula, and the Algerian–Tunisian coast. These regions have recorded more than twelve cold pools per year for a 10-year period (Llasat and Puigcerver 1990). If a cut-off low is defined as a closed depression at altitude that has become isolated and completely detached from the circulation associated with the polar or subtropical stream (and moves independently of that flow), the role that cold pools play in some heavy rainfall episodes are encompassed within the cut-off low definition (in other words, bringing in warm, moist air at low levels) and the vorticity advection. Adverse Weather Phenomena (AWP) are considered to be induced by Mediterranean cyclones when they are recorded at a distance less than 600 km from the cyclone-affected area. Following the WMO (1995) proposal, a meteorological phenomena should be considered an adverse weather phenomena if its intensity, frequency, or extension departs markedly from normal climatic conditions. In the case of high rainfall or wind storms, they become adverse when their intensity exceeds certain thresholds. Thus, it is possible to distinguish between: r Extraordinary: A precipitation input very much

above normal.

r Dangerous: A daily rainfall amount or rainfall dura-

tion period very much above normal. A stormy wind. r Catastrophic: An absolute daily rainfall amount or storm extent very much above normal. Rainfall duration or storm extent very much above normal. The absolute speed of a wind storm above 17.2 ms−1 or covering a wide area. This classification is based on the potential level of destructiveness of the particular adverse weather

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phenomena (Radinovic 1997). The categories are a function of nature, intensity, and spatial extent. The AWP that are not destructive, but do influence human life and activities are the extraordinary ones. Dangerous weather phenomena are considered to be those

which directly affect human life and material goods. Catastrophic weather phenomena include the dangerous phenomena that affect an extremely large area or reach the absolute maximum of intensity at some station or at some grid point.

Storms and Floods

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Fig. 18.2. The mean annual frequency of cyclones in the Mediterranean in the summer (upper map) and in the winter (lower map). Cyclone centres have been mapped in a grid of 2◦ latitude by 2◦ longitude for the period June 1998 to May 2001 (modified from Gil et al. 2002).

Jansà et al. (1996) have identified the presence of a cyclonic centre in 79 per cent of 721 heavy rainfall events in the Mediterranean region. Recently, Rigo and Llasat (2003) have shown that 79 per cent of heavy rainfall events occurring in Catalonia, Spain, were related to a cyclone. In the vast majority of cases, Mediterranean depressions form on the coast of Algeria due to the interaction of the wind with the Atlas Mountain chain. In cases where the rains affect Catalonia, Languedoc-Rousillon, or Provence-Alpes-Côte d’Azur (Figure 18.1a), the depression then shifts northwards. In order to lift the mass of warm, moist air there has to

be a forcing mechanism, which can be either meteorological (e.g. a convergence line, front, dry lines, etc.) or topographical (e.g. the coastline or, more usually, the mountains). It has been observed that the position of the depression plays a decisive role when it comes to localizing the zone in which the storms will occur, especially in the case of the southern Mediterranean. The analysis of some recent episodes has shown how the storms can shift with the movement of the depression (Llasat et al. 2003a). Contrary to the common view that is mainly promoted by mass media, the presence of a cold pool or cut-off

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low is not a necessary condition, nor is it sufficient on its own, for the occurrence of heavy rain. When cutoff lows are associated with such events, their role is to create a rather deep cyclonic circulation in the lower troposphere so that Atlantic air may be carried over the Mediterranean Sea. If the upper vortex is in the right location, this air, which is warm and very humid in the lower layers, impinges almost at right angles on the coastal mountain ranges and the forced ascent is enough to trigger the potential instability. Rainfall amounts over 200 mm in 24 hours are not uncommon in such cases, particularly on the eastern Spanish coast, where 40 per cent of catastrophic floods are associated with the presence of a cold pool. Due to the confusion generated by incorrect use of the term ‘cold pool’, some Meteorological Offices (Martín León 2003) have recently suggested using the term DANA (Depresión Aislada en Niveles Altos, Isolated Depression at High Levels) for those that could be identified as a cut-off low or COL (Llasat et al. 2007).

Convective Systems and Heavy Rainfall Events in the Mediterranean Region Beginning with the initial classification made by Byers and Braham (1949), and incorporating the most recent classifications (Doswell et al. 1996; Rigo and Llasat 2004), along with data from meteorological MCS-TS

radar, heavy rainfall events recorded in Mediterranean areas can be associated with the following structures (Figure 18.3): 1. Mesoscale Convective System (MCS): a precipitation structure may be identified as an MCS when its major axis has a length equal to or exceeding 100 km for three hours or more, and a minimum of 30 per cent of its area in each radar image can be associated with convective rainfall. An MCS is classified depending on the position of its stratiform region in relation to the convective area, and depending on the organization of the convective region. Thus, they could be divided into wellorganized systems (usually organized lineally) and poorly-organized systems, or clusters of convective structures (CLU). The first type could be divided into TS (with trailing stratiform area), LS (leading stratiform region), and NS (with practically no stratiform precipitation). A specific case of the NS class would be when the stratiform region is located on a flank of the convective line and its movement is parallel to the movement of the convective region. 2. Multicell systems (MUL): if a minimum of 30 per cent of the area covered by the precipitation structure in each radar image can be associated with convective rainfall, but does not meet the time and size conditions of an MCS.

MCS-LS

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Fig. 18.3. Classification of cloud systems associated with heavy rainfall events in the Mediterranean area (after Rigo and Llasat 2004). Key: MCS (mesoscale convective system), CLU (clusters of convective structures), TS (with trailing stratiform area), LS (with leading stratiform regions), NS (with practically no stratiform precipitation), MUL (multicell systems), IND (isolated convection), EST-EMB (convection embedded within stratiform rainfall), and EST (stratiform).

Storms and Floods

3. Isolated convection (IND): when small-scale, independent, and separate convective structures are identified. 4. Convection embedded in stratiform rainfall (ESTEMB): a stratiform region with some convective nuclei. The area covered by the convective precipitation does not exceed 30 per cent of the total area covered by the precipitation structure. 5. Stratiform (EST): when convective precipitation does not exist or does not exceed 3 per cent of the total area covered by the precipitation structure. 6. Supercell structures: if the precipitation structure has one large convective nucleus and a mesocyclone is detected within it (Doswell and Burgess 1993), a feature only detectable by using Doppler wind information. At present, the only systematic analysis of convective systems and heavy rainfall events made in the Mediterranean region has been done for the north-east of Spain (Rigo and Llasat 2004) (Figure 18.1). This analysis shows a total of 167 main precipitation structures associated with thirty-one heavy rainfall events recorded between 1996 and 2000. MCS and Multicells are the most common structures, with fifty-seven cases in each category, followed by individual cells (thirty-two cases). The cluster structure is the most frequent among the MCS cases (more than 50 per cent). Although it is possible to record all the structures in any season of the year, they arise mainly during the autumn season. An exception has been identified for the stratiform structures, which are not recorded in either spring or autumn when heavy rainfall events are produced. In the same region, an analysis of the longest rainfallrate series in Europe (1927–81) shows that 92 per cent of rainfall episodes in Barcelona (Spain) are nonconvective, yielding 63.5 per cent of the total precipitation, while 8 per cent are convective and provide 36.5 per cent of the total precipitation (Llasat 2001). To break this proportion down further, 3.8 per cent pertain to slightly convective episodes (less than 25 per cent of the total precipitation has a 5-min intensity of 35 mm per h), 2.9 per cent to moderately convective episodes (between 25 per cent and 75 per cent of the total precipitation above the previous threshold), and 1.3 per cent to strongly convective episodes (more than 75 per cent of the total precipitation exceeds 35 mm per hour). The intensity and duration ranges established by the above classification coincide broadly with the proper-

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ties of the various convective systems. The episodes termed strongly convective, of short duration and high average intensity, would have their origin in unicellular or multicellular storms (the life-cycle of a typical storm cell ranges between 20 and 30 minutes), arising mainly in summer. The moderately convective episodes that are of high mean intensity—with a duration exceeding a few hours—would correspond to highly organized convective systems composed either of multicellular or Mesoscale Convective Systems in which it is possible for 50 per cent of the precipitation to be stratiform.

A General Conceptual Meteorological Model Associated with Heavy Rainfall in the Mediterranean Region Synoptic and mesoscale factors responsible for heavy rainfall are not always the same either within a region or in different regions. It is nevertheless possible to speak of a basic conceptual model associated with those heavy rainfall events that produce catastrophic floods in the north-western Mediterranean region. This conceptual model (Llasat and Puigcerver 1992; Jansà et al. 1995, 1996; Palmieri and Clericci 1992) shows a long anticyclonic situation over the Mediterranean—for the days leading up to the events—which favours the formation of a Mediterranean air mass (Figure 18.4). Normally, the presence of any Mediterranean low or a convergence line organizes the differentiated air currents as well as internal low frontal boundaries. The intersection between the tip of a warm-wet current and a thermal-humidity boundary is the most likely place for reaching or releasing the convective instability and for the development of large convective clouds producing heavy rain. If the situation remains more or less stationary, the accumulated rainfall can reach very large amounts. Another possibility is to substitute the low internal boundary with a mountain barrier that stops and forces the ascent of the warm-wet current. Usually, both variants appear in combination and it is difficult to assess the relative contribution of each factor, bearing in mind the synergetic effects between them. As Figure 18.4 shows, the presence of the anticyclone over the central Europe–north Mediterranean area encourages the entry of a low-level flow from the south or south-east over the region, as well as the evaporation and accumulation of water vapour over the western Mediterranean. The Mediterranean is a semi-enclosed

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Storms and Floods

sea of no great depth, so it forms a large source of heat and moisture. Moreover, in many recent storms it has been observed that the temperature of the sea is even higher than usual and it is now well established that heat transfer and thermal convection reinforce cyclogenetic processes in the Mediterranean region (Alonso-Sarría et al. 2002). Typically, there is an Atlantic depression situated in the western part of the Iberian Peninsula or France (with an associated front that moves towards the north-east) and that helps to strengthen this flow. The position of the anticyclone and/or low is essential in organizing the flow that will affect one region or another. Romero et al. (1999) have shown that the majority of torrential rainfall events produced in Mediterranean Spain (Chapter 8) are associated with the presence of closed cyclonic circulations or shortwave troughs located to the south or west of the affected area. Good examples of this type of situation affecting Italy are two different events which impacted the Tanaro River basin in north-west Italy on 14–18 November 2002 and 24–6 November 2002. The synoptic scenario involved a sequence of two deep troughs progressively extending their influence over the Mediterranean basin. In both cases the flow became southerly and moist air impinged against the Alps and Apennines generating heavy precipitation. A further deepening of the second trough associated with a cut-off event took place between 25– 6 November. As a result, a deep depression was isolated to the south of Sardinia and produced wet south-easterly flow from the Tyrrhenian Sea towards the western Alps and Apennines (Taramasso et al. 2005). It is quite common to find moving systems that, in the course of their journey, affect different Mediterranean regions or countries. In Spain, for example, a convective system can start in the south and move towards the north-east, affecting all the regions along its path. This was the case in the flooding of 19 October 1982 which claimed thirty-nine victims in the Levante region of Spain, with over 500 mm of rainfall in 24 hours. This same system had previously produced rainfall exceeding 100 mm in 24 hours in Andalucia and did the same in Catalonia. Similarly, heavy rainfall systems that affect both Catalonia and the south of France are not infrequent (Llasat and Puigcerver 1994). An example was the serious flooding that occurred during 6– 8 November 1982 and the regional pattern of rainfall for that event is shown in Figure 18.5. In addition to affecting the Eastern Pyrenees region (Figure 18.1b), where 610 mm of rainfall were recorded over the three days, 150 mm (more than double the monthly average) were recorded in southern Andalucia on 6 November. A similar situation had arisen in October 1940, with

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more than 800 mm in 24 hours in the Eastern Pyrenees (Llasat 1993). Similarly, the heavy rainfall events that occurred in October 1986 and 1987, and again in November 1988, with more than 200 mm in 24 hours, caused floods in both the French and Spanish zones of the Eastern Pyrenees (Llasat and Puigcerver 1992; Ramis et al. 1994, 1995). Floods that are caused by a single meteorological event can affect Spain, France, and Italy and this was the case in September 1992. The first heavy rainfall events were recorded in the south of Catalonia and in Levante (eastern Spain) on 26 September, with total quantities exceeding 100 mm in many places. In the Rousillon region of France, the maximum rainfall recorded was 324 mm, 93 mm of that in 1 hour. On 27 September the rains mainly affected the Liguria region of Italy, where a total of 459 mm was recorded in 24 hours, with a maximum mean hourly intensity of 75 mm. On 28 September the rains mainly affected the Tuscany region of Italy, with a maximum of 148 mm. References to the direct role played by orography in heavy rainfall events are frequent. In the case of the 26–8 September 1992 event, the coastal mountains (over which the wind impinged perpendicularly) triggered the convection which remained latent over the Mediterranean Sea, with high CAPE (Convective Available Potential Energy) values, strong quasi-geostrophic vertical forcing, water vapour nourishment, and instability at low and medium altitudes (Llasat et al. 1999). In the Versilia event (Tuscany, Italy), recorded on 19 June 1996, heavy precipitation over the course of twelve hours occurred (with a maximum accumulated value of 478 mm) as a consequence of strong potential instability with large CAPE values and a low-level jet mainly driven by the local topography (Frontero et al. 1997). Sometimes the focusing of the convection and the quasi-stationary nature of the system can be explained by taking into account the relationship between the mountains and the dominant flow at low levels (Riosalido et al. 1997). On those occasions, the slow translation of the system is a consequence of the weak wind and the formation of new cells in the opposite direction to the movement of individual cells, in conjunction with winds blowing perpendicularly to the mountain ranges which are more than 2,000 m above sea level. This was the case for the flash flood that occurred in Biescas, Huesca (Spain) in the Central Pyrenees, on 7 August 1996. The flood completely destroyed a campsite that was situated between two small gullies that overflowed. Out of a total of 630 people registered at the campsite, 183 sustained

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injuries and a further 87 died in less than 45 minutes. The maximum precipitation was 269 mm in 24 hours, with 226 mm in only 3 hours and 153 mm between 16 and 17 hours UTC. However, orography can also play an indirect role in heavy rainfall development. Between 28 September and 5 October 1987, 431 mm of rain was recorded in Barcelona (Catalonia, Spain). On that occasion the main role of the Pyrenees was related to the generation of a strong orographic dipole, giving rise to a strong mesohigh over Catalonia. As a consequence of this high, the synoptic south-easterly wind at low levels was substituted near the Catalan coast by a westerly wind and a convergence line developed over the Mediterranean Sea. The situation remained quasi-stationary for a few days and the continuous vertical forcing of the convective and unstable wet air gave rise to the generation of an MCS (Ramis et al. 1994). In contrast, in eastern Mediterranean countries where heavy rainfall events with more than 100 mm in 24

hours are rare phenomena, the role of orographic factors is secondary or non-existent in rainfall production (although it can play an important role in cyclogenesis). For example, an extreme rainfall event (maximum input of 44 mm per day in the Balkans and Asia Minor) over the Black Sea, Ukraine, and Russia occurred from March 28 to 1 April 1995. The event started in the Mediterranean, when a secondary cyclonic eddy was formed in a deep trough over northern Italy and moved over the Adriatic Sea, although the main factor was a frontal system driving the circulation (Chakina and Berkovich 1997). In Greece, many floods are produced by heavy but short rainfall events, and the role of deforestation and urbanization is often very important in their genesis. There is some evidence to suggest that they mainly affect urban areas, due to the transformation of urban catchments and river channels. They are more destructive in the western parts of Greece due to climatic, geomorphological, and vegetation factors and the impact of human

Storms and Floods

activity. The most serious flood of the last ten years in Greece was in 1994, when a thunderstorm produced 68 mm of rain in one hour in Athens. Unlike the western part of the basin, in the eastern part of the Mediterranean, the main flood events commonly occur in spring.

Classification of Floods and Storms in the Mediterranean Region Floods are a complex hydrometeorological hazard. Meteorological and climatic factors, drainage basin factors, drainage network and channel morphometrics, and human factors play a major role (López-Bermúdez et al. 2002). Heavy rains, long rainy periods, or snowmelt are necessary but not sufficient to cause them. Other factors such as antecedent precipitation, terrain, and surface run-off characteristics are also important (Chapters 6 and 8). Natural processes also interact with human activities. For example, land use and its history and the civil and hydraulic infrastructure can have very variable effects on the natural patterns of flood generation and impact. It is not easy, therefore, to create a Mediterranean flood classification that can be applied to all catchments and events. The classification presented here attempts to integrate some of the meteorological, hydrological, and impact aspects of Mediterranean floods, although hydraulic aspects have not been considered. In order to understand the various types of flood events that can affect Mediterranean countries, it is important to remember from the outset that Mediterranean river systems comprise three basic types: r Type A: High-mountain rivers (>2,000 m) with

rainfall- and snowmelt-influenced regimes with large basins (>2,000 km2 ) and perennial flows. r Type B: Rivers with upland headwaters with intermediate-sized catchments (50–2,000 km2 ) and flows. Due to the highly seasonal rainfall regime (with an extended dry season) and the abstraction of river flows, some have ephemeral reaches in their catchments. r Type C: Short littoral water courses with small catchments (5–50 km2 ) and steep channels. They are ephemeral systems marked by infrequent but torrential flows. A key control on the behaviour of these rivers is the torrential rainfall events discussed in the first part of this chapter and inherent to the Mediterranean region. Accordingly, a distinction can be made between four kinds of floods and associated flood-generating mech-

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anisms in Mediterranean catchments, and these are described in turn below.

Type 1: Short-lived, High-intensity Rainfall Events The first type of storm involves short-lived events (less than 3 hours and often less than 1 hour) of very intense precipitation (peaks of rainfall intensity above 3 mm per minute) but limited overall rainfall totals (2000 km2 ). This kind of event requires convective instability with abundant feeding of warm and wet air from low levels, and a mechanism to force air ascent to release the potential instability or to destabilize the air column. Convective rainfall is generally produced by ‘multicells’ or ‘mesoscale convective systems’. It is possible to distinguish between two kinds of flood event in this category. Type 2a lasts less than 24 hours and the maximum precipitation is usually recorded in less than 6 hours, with accumulated rainfall of nearly 200 mm. They are ‘strongly convective events’ and can produce catastrophic flash floods in Type B and Type C rivers that are simultaneous with the maximum rainfall input. Type 2b events last more than 24 hours but generally less than four days. Although accumulated rainfall usually has values between 200 and 400 mm, values of more than 800 mm are possible. Peaks of strong rainfall intensity and moderate but continuous rainfall are recorded successively. Consequently, they are ‘moderate convective events’, and while they can also produce local flash floods in Type C rivers the floods occurring in Type B river catchments (and occasionally Type A rivers also) are the most significant. Floods of Type 2a have produced the highest number of casualties when they have affected flood-prone areas with high concentrations of people. Damage produced by catastrophic floods of Type 2b relates to total or partial destruction of infrastructure (e.g. houses, bridges, and roads), power cuts, urban inundation, agricultural and livestock losses, and, frequently, loss of human life. These events usually occur in autumn, although some cases have also been recorded in spring or summer. Figure 18.7 shows the rainfall map and the changes in rainfall intensity associated with the catastrophic flood that affected Catalonia on 10 June 2000. The maximum cumulative precipitation was 224 mm and this produced a peak flow of 628 m3 s−1 in the ephemeral channel of the Riera de la Magarola (a 96 km2 tributary basin of the Llobregat River) that caused extensive damage (Figure 18.8). The peak discharge of the main channel of the Llobregat River at its mouth was estimated at 1300 m3 s−1 . The maximum recorded peak flow for the Llobregat River of 3080 m3 s−1 was reached during the Type 2b floods of 20–3 September 1971. Another example of a Type 2a event was the storm of October 1988 that produced a flash flood on the Cadereaux River in southern France. On 3 October the flood peak reached 600 m3 s−1 in a catchment area of only 42 km2 . The flood destroyed many cars and some buildings in the city

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of Nîmes in southern France. The event was produced by a rainfall input of 300 mm in six hours (Stanescu, personal communication). The flash flooding produced in southern France on 8 September 2002, with 680 mm of rain in 24 hours, was also related to a Type 2a event (Milelli et al. 2006). During this event an area of 5,393 km2 recorded more than 200 mm of rainfall and large areas were inundated across the Gard, Herault, and Vaucluse departments in southern France (Ruin et al. 2008). Figure 18.9a shows extensive flooding in the village of Aaramon near Nîmes and Figure 18.9b shows flood waters of the River Crieulon over-topping a dam near Sommières. In Liguria in north-west Italy on 13 August 1935, a precipitation event of 389 mm in eight hours produced flash flooding along the Orba River (141 km2 ) with a peak flow of 2200 m3 s−1 . This is also a clear example of Type 2a event. The floods recorded during 14–18 and 24–6 November 2002, with maximum accumulated rainfall in 72 hours of 560 mm and 442 mm respectively, could also be considered as Type 2b events. During September 1992, two catastrophic Type 2b flood events occurred over the south of France and the north of Italy. The first took place between 22 and 23 September and the peak flow of the flash flood produced on the Ouvèze River (580 km2 ) in Vaisson-la-Romaine (PACA), was above 1000 m3 s−1 . The second event started on 26 and ended on 28 September in central Italy. Another catastrophic flood event of Type 2b was recorded during 12–13 November 1999 in the Aude region (LanguedocRoussillon), with 624 mm in 36 hours.

Type 3: Long Duration, Low-intensity Rainfall These are episodes of long duration (approximately one week) with relatively weak average rainfall intensity values, although there may be peaks of high intensity within the overall distribution. Total precipitation during Type 3 events can be >200 mm. If floods occur, they are usually in larger rivers of Type A and B as described above. They may be described as ‘slight convective events’ and are usually associated with convection embedded in stratiform rainfall (Figure 18.3). Although not very frequent, they usually occur in winter and, occasionally, in spring. The floods recorded in Catalonia in January 1996 (Llasat et al. 2000) are a good example (Figure 18.10). River catchments on the Balkan Peninsula have also seen Type 3 events. The city of Celje in Slovenia, for example, that lies on the confluence between the Savinja (1192 km2 ) and Voglarinja

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part of the Mediterranean, such a pattern may be the result of depleted fuel loads related to socio-economic factors.

Fire Regime: Fire Size and Return Times A key factor contributing to the marked spatial and temporal variability of fires is the extreme skewness of the distribution of fire magnitude. Most of the burnt area

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Francisco Lloret, Josep Piñol, and Marc Castellnou 60 Number of fires Area burnt

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is accounted for by a very small proportion of fires. These major fires can achieve catastrophic dimensions in terms of their extent and intensity as witnessed, for example, in the Peloponnese region of Greece in the summer of 2007. This can be clearly seen, for example, using the fire record of Mediterranean France (Figure 19.3). The largest 0.15 per cent of fires (>1,000 ha) burnt 39 per cent of the total area, the largest 1.4 per cent of fires (>100 ha) burnt 73 per cent of the total area, and the largest 7.4 per cent of fires (>10 ha) burnt 90 per cent of the total area. This is an almost universal pattern for any region with wildfires—including non-Mediterranean environments—and it has been well documented for the whole of the Mediterranean basin (Ricotta et al. 2001). This pattern is of great importance for managers in order to establish strategies to avoid these extreme events, when uncontrolled wildfires cause the most impact on properties and natural resources and threaten human lives. In Portugal, two periods of extreme fire hazard in the first half of August and the second week of September 2003 resulted in simultaneous fires that are estimated to have burnt more than 440,000 ha and killed twenty-one people (European Commission 2004; Viegas 2004) (Figure 19.4). In July 1994, again during extreme conditions of high temperature and winds, seventeen fires of more than 1,000 ha—many of them occurring simultaneously—burned 170,000 ha

and produced twenty-two casualties in Spain (Moreno et al. 1998). The occurrence of simultaneous fires is an additional cause of increased fire size, as fire-extinction resources become overstretched. These extreme episodes also introduce a source of uncertainty in the temporal evolution of the burnt surface, as they can destroy any stabilization tendency. In Portugal, the annual burnt area fluctuated steadily around 50,000 to 150,000 ha prior to the extreme event of 2003 (European Commission 2004). As stated above, burnt area estimates are spatially variable because of the different methodologies used in each country (Vélez 2000). In many cases the analysis of fire origins and consequences relies on accurate mapping of the extent of the burnt area. Many research initiatives have shown the value of satellite-based remote sensing to obtain reliable and objective estimates of these areas across large territories (Salvador et al. 2000; Chuvieco et al. 2002). The spectral properties of the burnt areas determine the reliability of the available methodologies (Pereira et al. 1999). The visible spectral domain (e.g. from Landsat Thematic Mapper TM and National Oceanic and Atmospheric Administration’s Advanced Very High Resolution Radiometer NOAA/AVHRR) has been extensively used for mapping burnt areas. In this spectral domain, recently burnt areas are relatively dark, but they can be difficult to discriminate from other dark land covers such as water bodies and some soil types. However, burnt areas can be distinguished because they tend to produce post-fire signals resulting from vegetation recovery. In the near-infrared (NIR) spectral region, the signal of recent fires is strongest—particularly when large amounts of charcoal are produced—and it has been widely used for fire mapping. A common limitation of these methodologies is that only fires above a given threshold (30–100 ha) can be accurately detected and mapped. However, since most of the burnt area is the result of large fires, the unaccounted for surface is a small proportion of the total and can be estimated by regression (European Commission 2004). Since 2001 the European Commission routinely provides annual estimates of the area burnt in fires larger than 50 ha in Portugal, Spain, France, Italy, and Greece based on WiFS or MODIS imagery (European Forest Fire Damage Assessment System, EFFDAS; European Commission 2004; Figure 19.4). This information may be used to establish several parameters of the fire regime related to the temporal sequence of fires. The interval between fires is the elapsed time between two consecutive fires in a geographical location. Fire period is the length of time needed for a given area to experience a fire; it is the inverse of fire

Wildfires

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Fig. 19.4. The area burnt by forest fires in Portugal during 2003. Burnt areas are shown in black over a grey-composite of satellite images (European Commission 2004). The 2003 fire season in Portugal was the worst ever recorded, with over 400,000 hectares burnt, a figure more than double that of the previous highest year (1991). There were twenty-one casualties in 2003 (Viegas 2004).

frequency, that is, the number of fires occurring in a given area per unit of time. Unfortunately, the length of the survey period is generally too short to establish such parameters accurately, particularly when temporal and spatial variability is high. Some researchers have found that fires are not randomly distributed, but

that vegetation, climate, topography, and human activity determine the spatial pattern of fires (Vázquez and Moreno 2001; Díaz-Delgado et al. 2004a ). These findings agree with rough estimates obtained from administrative sources and offer the advantage of covering larger regions (Moreno et al. 1998).

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Díaz-Delgado et al. (2004b) used Landsat MSS images to produce yearly maps of fire occurrence using the NDVI (Normalized Difference Vegetation Index) to describe the fire regime in Catalonia (Spain) from 1975 to 1998. They found that the estimated mean interval of fire occurrence in landscape units of c.8 km2 was 5.5 years with a fire period of 133 years, a value similar to the 100 years estimated for the whole of Spain from administrative records (Moreno et al. 1998). Fire hazard appears to increase asymptotically before reaching a plateau at around 15–16 years after a fire. However, it is important to stress that recurrent fires occur quite often even in recently burnt areas, since 14 per cent of the burnt area experienced at least one additional fire during the 24-year period, with a maximum of up to five fires in areas with strong winds.

Historical Context The Holocene Since the physical determinants of fire (fuel, heat, and oxygen) are basically independent of human activity, provided natural sources of ignition occur, wildfires should have existed before human settlement. This seems to be the case for regions such as California, Australia, and South Africa where lightning storms are not uncommon and wildfires have occurred since prehistoric times (Kemp 1981; Keeley 1992; Keeley et al. 1999). Lightning storms also occur in the Mediterranean basin (Vázquez and Moreno 1998), although the surface area burnt by this kind of fire is now commonly much reduced compared with earlier periods of the Holocene. Lightning fires usually start under moist conditions—they therefore behave as low-intensity fires that, in the past, could have continued until the onset of heavy rain. Records of charcoal particles from lake sediment sequences (often conducted in association with pollenbased reconstructions of vegetation change) constitute an important indicator of past regional fire occurrence (MacDonald et al. 1991). Parts of the Mediterranean have well-preserved, high-resolution lake sediment records (Chapter 9) and some good records of Holocene fire history are beginning to emerge (e.g. Sadori and Giardini 2007). In these studies, the charcoal counts often provide data on large, high-intensity fires, with much less information on low-intensity fires. Most of the available information comes from the western part of the Mediterranean basin, showing that fires were occasionally present in the Iberian Peninsula during the last cold stage (c.30,000 to 15,000 BP) (Carrión 2002).

However, a significant increase in charcoal deposition does not arise until the establishment of much warmer and seasonally dry conditions, typical of the Mediterranean climate (Jalut et al. 2000). The precise dates may be different between localities, but from c.6,000 BP there is a clear increase in charcoal in such records. This pattern is consistent in several localities in the southern (from c.4,200 BP in Sierra de Gador) (Carrión et al. 2003), south-eastern (from c.5,000 BP in Sierra de Segura) (Carrión 2002), and the central Iberian Peninsula (from c.5,000 BP) (Franco-Múgica et al. 2005). In fact, the use of fire by people as a driver of landscape transformation is documented in the mid-Holocene in French sub-mediterranean mountains, where Pinus sylvestris forests moved to the present-day rangeland during the period 4,800–3,000 BP (Quilès et al. 2002). In all these records, fire occurrence increases up to the present day, believed to be largely as a result of human activities. This trend is well illustrated by the charcoal peaks obtained in the eastern Iberian Peninsula corresponding to periods of intense landscape transformation during the Roman agricultural expansion (c.2,000 BP), the medieval increase in grazing area (c.1,500–1,300 and 1,000 BP), and, more recently, after the eighteenth-century population rise (Riera-Mora and Esteban-Amat 1994). Overall, although evidence for fires is present, it seems that biomass burning was low in the western Mediterranean basin during the early Holocene (Carcaillet et al. 2002) and that human activity has been a major driver of fires since about 2,000 BP. Data from other parts of the world with Mediterranean-type climates suggest that fire has long been a major factor in shaping these Mediterranean ecosystems. In south-western Australia, for example, high-resolution records from the Pliocene have shown the existence of fires at intervals of 5–13 years prior to the establishment of human populations (Atahan et al. 2004). In the same region, there is evidence for fire throughout the Holocene, but with regional and temporal variability (Dodson 2001). In southeastern Australia, charcoal has been reported to be strongly associated with sclerophyllous vegetation during the last interglacial. Once established, the fire regime demonstrates a reasonably constant pattern of occurrence even during the last glacial period and the Holocene, suggesting that humans may have contributed to this regular pattern of fire occurrence (Singh et al. 1981). In California, by contrast, charcoal deposition indicates that the regime of large fires has not changed dramatically over the last 500 years, showing a strong relationship with climate. Here they can be assumed to be part of the natural fire regime of

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the region, where smaller fires would also be present (Mensing et al. 1999). Overall, it appears that in recent millennia, in these other Mediterranean regions, fire regime has not experienced such strong fluctuations over the fine temporal scales that are evident in the Mediterranean basin, where historical events have produced a changing pattern of fire occurrence. Fire recurrence intervals in these regions are estimated at 4–45 years in the South African fynboss (van Wilgen et al. 1992), 30–60 years in the Californian chaparral (Keeley 1992), and 3–13 years in the Australian heath lands (Atahan et al. 2004).

Historical Times and Data Sources For the historical period features such as fire scars on trees can provide information on low- or mediumintensity fires (Figure 19.5). Although this information is often very scarce, in the north-east Iberian Peninsula low-intensity fires have been documented at a frequency of around twenty-five years since the beginning of the eighteenth century (Pellisa, personal communication).

Fig. 19.5. A fire scar at the base of a pine tree trunk in Catalonia, north-east Spain. This feature is indicative of a low-intensity fire that did not burn the forest canopy (photo: Marc Santandreu).

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This frequency has permitted the persistence of a mature structure of Pinus nigra forests by allowing tree survival but at the same time thinning the dense regeneration (Castellnou et al. 2002a ). Written sources of information are also available, including occasional news of fires, or even detailed records. This is the case in the archives of the city of Tortosa in Catalonia for the fourteenth and fifteenth centuries. The expenses of the city were recorded in detailed accounts, including the payments to people who fought fires in the surrounding forests. It is possible, therefore, to reconstruct the frequency of fires during this period. Although the comparison between historical periods has some significant shortcomings due to differences in the accuracy of the recorded information, the results suggest the occurrence of a number of fires of the same order of magnitude as in the last decades of the twentieth century (Lloret and Marí 2001). The size distribution of fires— with many small fires and very large ones—is also comparable between these periods. Occasional information from travellers provides only fragmentary data that do not allow the importance of fires to be established. These observations often show a scenario where burning was quite common for improving pastures, while in other cases the records describe extraordinary events. Legislation that banned such fires in forested areas is often found in several areas of the Mediterranean (including Sardinia, Catalonia, Castile, and Crete) (Grove and Rackham 2001) that were associated with the exploitation of forest goods. In summary, it appears that two different attitudes to fire coexisted in many parts of the Mediterranean until the twentieth century. One approach involved the use of deliberate fire ignition to obtain forest clearings and managed pastures, while the other was based on fire prevention by banning burning activities and prosecuting those who carried them out. In some cases it resulted in active fire-fighting. Landownership and historical circumstances would control the balance of these activities. These, combined with climate conditions, would determine the historical pattern of fire occurrence in a given area. Although a detailed reconstruction of fire history is not available for any country, it seems that there was not a regular pattern of fires across time. Instead, periods of variable burning incidence took place prior to a steep increase in the last two centuries (Riera-Mora and Esteban-Amat 1994). In the late nineteenth and early twentieth centuries, administrative and journal sources provide more regular sources of written information. From Spanish administrative records, we know that at the end of the nineteenth century, fire was frequent in Spanish forests,

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with a similar frequency and size structure (most of the surface is burned by a few fires) to the present (Araque 1999). Some localities were particularly prone to burning, probably due to fire ignited by traditional rural communities. In fact, in this period, ideas on fire prevention and fire-fighting were introduced by forest engineers in a rural society where fire was commonly used. While catastrophic fires occasionally occurred at the beginning of the twentieth century, this conflict remained, and eventually fire-fighting strategies became dominant.

Determinants of Fire Regime: Climate, Vegetation, and Social Context Combustion is an exothermic reaction determined by three main factors: fuel, heat, and oxygen. Although the chemical processes involved in fire are relatively simple, the controls on fire regime are rather more complex, since historical and biological processes are also involved, as shown in Figure 19.6. Fuel provided by Mediterranean vegetation is particularly prone to produce high-intensity and fast-spreading fires. Mediterranean plants usually accumulate dead branches and leaves in standing shoots and on the ground. Low decomposition rates associated with water scarcity during long periods enhance fuel accumulation. In addition, this material shows a high surface area to volume ratio, which is a consequence of the reduction of the evaporative surface area on the leaves of Mediterranean plants, and of the common occurrence of defensive structures in the bark, shoots, and leaves (Chapter 7). Such a high ratio has two consequences. First, water content largely fluctuates according to drought periods. As the water content decreases in fuel, less heat is needed to propagate the combustion reaction. Second, the chemical reaction is facilitated because more oxygen per surface unit is in contact with fuels. The chemical characteristics of Mediterranean vegetation also contribute to promote the spread of fire, as volatile compounds are abundant. These compounds, mainly terpenes, are extremely flammable and have relatively low boiling points, thus further promoting the spread of the burning front. The Mediterranean climate is characterized by strong seasonality, with high temperatures and low air humidity in summer (Chapter 3). At this time of year the water content of fuel is very low, which facilitates the ignition and propagation of fire. Also, local winds increase both fuel desiccation and fire propagation, creating

favourable conditions for large and intense fire events. Accordingly, climatic records from eastern Spain show that in wet summers the area burned is lower than in dry summers (Pausas 2004). However, climatic conditions do not remain constant throughout the whole summer. In fact, a wildfire is a stochastic event that is usually concentrated within a few days. This may produce difficulties for attempts to correlate seasonal or annual average climatic parameters with fire occurrence in terms of the number of ignitions and amount of burnt area. Accordingly, the yearly number of days with extreme climatic conditions correlates well with the annual number and area of fires (Piñol et al. 1998). Climatic determinants of fire may, however, be rather more complex. The growth of vegetation, such as herbs and grasses, enhanced by abundant rainfall, may result in more fuel being available to burn in future summers. Indeed, Pausas (2004) has found a positive correlation between summer rainfall and summer burned area for a time lag of two years in eastern Spain. A particular challenge is to assess the significance of very recent climatic changes on fire regime. Interestingly, the recent rise in temperatures (about 0.35 and 0.27◦ C for annual and summer temperatures per decade respectively in eastern Spain) does not correlate with fire occurrence (Pausas 2004). However, the combined effect of rainfall and temperature produces a different pattern. Hence, in the same region the number of summer days with high climatic fire risk (that is, combining high temperatures and low humidity) has increased significantly during the twentieth century correlating with the rise in both the number and extent of fires (Piñol et al. 1998; Figure 19.7). The combination of natural factors promoting fire occurrence has been enhanced by human activity. Apart from the effect of human activity on climate change, humans directly modify the fire regime by creating ignitions or by fighting fires. The natural source of ignition is lightning. Although geographically variable, it accounts for about 10 per cent of wildfires (Vélez 2000). Lightning-caused fires contrast with human-caused fires, since the latter are clustered in certain geographical areas, concentrated in summer, and show smaller maximum sizes (Vázquez and Moreno 1998). Humans introduced fire as a land management tool in the Mediterranean region thousands of years ago when fire was deliberately used to improve grazing or to facilitate clearing. Nowadays most such practices are declining in the northern part of the basin, but rural traditions continue in large parts of Mediterranean North Africa. However, agricultural practices remain a major source of ignitions, even in

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Fig. 19.6. A conceptual model of factors influencing the area burnt in a region showing (a) without interaction between the factors, and (b) with the main interactions that influence fuel build-up.

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Fig. 19.7. (a) Changes in summer climatic fire risk in Catalonia (north-east Spain) between 1940 and 2000 (r = 0.57, p < 0.01), and (b) the burnt area in relation to the number of days with high climatic fire risk (exponential fit, r = 0.88, p < 0.01). Fire risk is estimated as the number of days exceeding a predetermined value of the climatic I87mod index. This index considers daily maximum temperature, minimum relative humidity, and maximum wind speed (see Piñol et al. 1998 for calculation details). (Redrawn from Piñol et al. 1998).

the European Mediterranean (Vélez 2000), while new sources have appeared, including those related to economic development (e.g. traffic, electrical power lines, rubbish dumping) and recreational uses. For instance, during the period 1968–98, most wildfires in eastern Spain started on Sundays (Terradas et al. 1998). Indeed a positive correlation between population density and number of fires has been observed (Terradas et al. 1998). A major effect of human activity in fire regime modification is related to socio-economic factors leading to profound changes in land use patterns (Moreira et al. 2001; Tàbara et al. 2003). The abandonment of traditional farming in many countries of the northern Mediterranean basin has taken place during the course of the twentieth century. Consequently, fuel load has grown as a result of vegetation succession. Paradoxically, this trend is enhanced by those fire-fighting policies that prevent the fuel reduction effect produced by small fires. The increase of wildfires in the last decades of the twentieth century is the logical result of the combination of huge fuel loads, increasing climatic risk, and multiple ignition sources. It is often advocated that fuel accumulation may be minimized by grazing or clearing. These were common practices a few decades ago, but they are no longer economically viable in most areas. Public funds to support these activities over vast areas are not feasible. By contrast, in North Africa, local populations keep fuel levels low by grazing or by collecting it for heating—as was common practice decades ago in the European Mediterranean countries. While wildfires seem to threaten the fate of forests in the northern Mediterranean, overexploitation tends to limit forest recovery in the south.

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Wildfire Hazard and the Impacts on Ecosystems Fire Hazard Fire hazard or risk has a structural component and a time-varying component. The structural risk arises from topography, vegetation characteristics, and some human-related factors (Chuvieco et al. 1999a). The time-varying risk comprises the relatively slow change in the water content of fuels and a much faster change in meteorological conditions such as wind direction and speed and air temperature and humidity. Structural risk is usually known from topographic and vegetation maps. Of these, the most useful are those that classify the vegetation within the so-called fuel models (Rothermel 1983). These classify vegetation types according to their most important characteristics that affect the spread of fire. Fuel model maps are usually obtained from a combination of field survey and photointerpretation, but the process can be speeded up by using satellite-based imagery (Riaño et al. 2002). Unless an important fire occurs, these maps only need to be updated every few years. The fuel moisture content (FMC) is estimated differently for dead and live fuels. In general, dead FMC is estimated from meteorological variables (Viney 1991), as it is the result of a simple physical process, that is, the exchange of water between air and the dead leaf or twig. As the estimation of the FMC of living fuels is much more difficult, this variable is often measured on selected locations and species (Viegas et al. 1992; Viegas et al. 2001). When it is desirable to have a map of living FMC it is necessary to extrapolate the key measurements using some kind of model based on variables routinely measured in meteorological stations (Viegas et al. 2001). Alternatively, living plus dead FMC can be estimated by satellite remote sensing, although the estimate is usually better for grasslands than for shrublands (Chuvieco et al. 1999). The most sensitive variables for FMC estimation are based on short-wave infrared bands, and the combination of vegetation indices and surface temperature (Chuvieco et al. 1999b). At an operational level, national fire agencies rely on the calculation of meteorological indices of fire danger. In the past each country used its own index, but a recent development has seen the Canadian Forest Fire Danger Rating System (van Wagner 1987) becoming the de facto standard fire danger index (Viegas et al. 1994). The Canadian system is not a single index, but a set of indices with different time-varying characteristics. Its different components are calculated from daily measurements

of air temperature and humidity, wind velocity, and precipitation. The indices can be predicted using forecasted weather variables. In practical terms, its component DC (Drought Code) was originally designed as a numerical rating of the slow-varying moisture of deep, compact organic layers. Over time, it also proved useful to forecast the fuel moisture of live fuels (Viegas et al. 2001). Aguado et al. (2003) showed a good relationship between spatially interpolated DC values and several satellite-derived indices of FMC. The component FWI (or Fire Weather Index—which is a numerical rating of fire intensity) is the most comprehensive indicator of the danger of fire. The European Commission provides an estimation of the fire risk forecast during the peak of the fire season, i.e. from May to October (European Forest Fire Risk Forecasting System or EFFRFS). The EFFRFS has been designed to compute several types of forest fire risk indices varying from short-term (including the Canadian FWI, Figure 19.8) to long-term indices. These long-term indices include the so-called fire probability and fire vulnerability indices. By 2004, this system delivered fire risk maps to a large number of EU Member States and to some Candidate Countries such as Turkey, Bulgaria, and Romania. Examples for 2001, 2002, 2003, and 2004 are shown in Figure 19.8.

The Impact on Ecosystems The most obvious impact of forest fires is the partial or total destruction of the vegetation. This impact has been traditionally measured as the area burnt. This parameter provides a simple estimation of fire impact that can be easily tracked over time and across countries, and it is assumed roughly to relate to both ecosystem and economic impacts. As described above, remote sensing has a lot to offer in burnt area mapping and provides a useful tool that has been combined with GIS techniques to analyse land use dynamics (Lloret et al. 2002), fuel distribution (Koutsias and Karteris 2003), or post-fire erosion hazard (Ruiz-Gallardo et al. 2004). The impact of fire on ecosystems depends on the characteristics of fire, especially its intensity, and this parameter shows great spatial variability, even within the edges of a single fire (Figure 19.9). In high-intensity fires most of the vegetation canopy is destroyed. However, the ability of Mediterranean vegetation to regenerate after fire is well known and this is due to several factors (Buhk et al. 2007; Chapter 7). Some species have structures that protect buds from extreme temperatures such as the bark of the cork oak or of Pinus pinea, and the leaf sheaths of some bunch grasses. Most commonly, many shrubs

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Fig. 19.8. Fire risk in Europe according to the Canadian FWI (Fire Weather Index) on 1 July of the indicated years—see text for explanation. Data are from the European Forest Fire Risk Forecasting System at the European Institute for Environment and Sustainability, CEC, JRC, Ispra, Italy; , accessed 5 December 2008.

(e.g. Arbutus unedo, Erica sp. pl.) are able to resprout from below-ground structures that are protected from the high temperatures by the soil. The soil also protects the seeds of some species (e.g. Cistus sp. pl.) that are able to germinate successfully and establish after fire. In

fact, fire itself releases a range of resources (e.g. light and nutrients) that favour vegetation growth. The germination of some species is even increased by fire as it helps to break the hard exterior of their seeds or to open fruits, for example Aleppo pine cones. Site productivity, largely

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Fig. 19.9. Burned forest in a highly populated area of Catalonia in north-east Spain. Note the spatial variability in fire intensity as indicated by the presence of areas with burned tree trunks and sites where the fire barely scorched the canopies (photo: Marc Santandreu).

due to water and nutrient availability, enhances postfire regeneration so that north-facing aspects, terraces, and valley bottoms, for example, show a greater development of vegetation (Pausas et al. 1999). On the other hand, low-intensity fires reduce the fuel load thereby decreasing the likelihood of future high-intensity fires. Their effects are not so dramatic and they allow the persistence of most tree populations and the soil seed bank. The ability of Mediterranean plant communities to regenerate after fire is not unlimited and high fire intensities reduce the ability of resprouting species to regenerate (Lloret and López-Soria 1993). Although high temperatures increase the germination rate of some species, when they achieve a threshold, seed viability declines (Herranz et al. 1999). As a result, the overall recovery of the community plant cover is negatively related to fire intensity (Díaz-Delgado et al. 2003). The season of fire occurrence may also determine the impact on the vegetation. If fire occurs before the replenishment of the seed soil bank, which for many species occurs in late spring or summer, germination would be scarce the following autumn (Dominguez et al. 2002). Resprouting may also be diminished if fire occurs when the level of stored resources in underground organs is low (winter and spring) (Cruz and Moreno 2001). Repeated fires at short intervals also strongly reduce post-fire regeneration. Seed banks cannot be restored if fire occurs before plants attain their reproductive age. As a result, pine populations may dramatically decline (Eugenio and Lloret 2004). Resprouting also decreases after recurrent fires because the bud bank (the set of meristematic tissues allowing the growth of new shoots) is eventually

depleted, and resources from underground organs are lowered before they can be restored from photosynthetic tissues (Canadell and López-Soria 1998). Overall, plant cover needs more time to recover after successive fires (Díaz-Delgado et al. 2002), and, over time, shifts in vegetation structure and composition may result (Pantis and Mardiris 1992). A recent review of the challenges faced by plants as they recolonize post-fire terrains has been provided by Buhk et al. (2007). Fire also affects some of the services that ecosystems provide, such as water resources and soil protection. In August 1990 a severe wildfire affected the Réal Collobrier experimental research basin (Var Department, France), where catchment hydrology had been monitored since 1966. Depending on the sub-basin, the proportion of burnt surface varied from 0 to 85 per cent. This fortuitous fire allowed a detailed study of the effect of fire on the hydrology of the catchments, as it allowed a direct comparison between affected catchments with controls, and the comparison of each catchment with its own pre-fire conditions (Lavabre et al. 1993). Among other findings, this study showed an increase of c.30 per cent in the annual runoff yield and an increase in flood frequency. The ten-year return period flow estimated before the fire was exceeded three times in a year, even though the rainfall events in that particular year were not unusual (Chapter 8). A post-fire increase in runoff can be considered universal (Belillas and Rodà 1993; Scott 1993), and it is simply the consequence of the reduction in evapotranspiration due to the destruction of the vegetation cover by fire. However, changes in hydrograph form and in flood frequency have not always been found following catchment burning. In a study of fire effects in several catchments in the Mediterranean climate zone of South Africa, Scott (1993) listed several conditions as indicators of a marked hydrological response after fire, and hence, of soil erosion: (1) a high level of soil heating, which itself is a function of fuel load, fuel and soil wetness, and weather conditions; (2) the loss of ground cover may limit vegetation growth and cause the site to be bare for longer after fire; (3) the roads and tracks may become extensions of the channel system after fire; (4) vegetation types that lead to the development of water repellent soils improve the chances of a sharp hydrological response to fire (in Scott’s study, this was high for pine plantations and low for the natural fynboss vegetation); (5) wildfires are expected to cause a stronger hydrological response and a higher erosion risk than less intense fires because they are more likely to occur in hot and dry conditions when soil and fuel moisture are low. Other studies in Italy and Spain have confirmed an

Wildfires

increase in soil erosion at the plot level after fires, and that the increase was more noticeable when the fire was more severe (Giovannini et al. 2001; de Luis et al. 2003). Recent research has also examined the impact of fires on air quality and on the flux of mercury from forest fires to the atmosphere. Cinnirella et al. (2008) used remotely sensed satellite data for 2006 to estimate that the total amount of mercury released to the atmosphere from this source in Mediterranean countries was 4.3 Mg year−1 .

Management: How can the Impact of Wildfires be Reduced? It has been shown that the area burnt increases with the number of ignitions, with the accumulation of fuel, and during adverse weather, and decreases with fire suppression (Figure 19.6a). However, there are interactions among these factors that complicate such a straightforward interpretation. A reduction in the number of ignitions or an increase in the suppression capacity would allow more fuel build-up that, in turn, will positively affect the area burnt (Figure 19.6b). As we have no control on the weather, there are, in principle, three possible ways to fight wildfires: 1. by reducing ignitions; 2. by increasing fire suppression; 3. by reducing fuel loads. The dominant fire-management policy in Mediterranean Europe is the suppression of all wildfires. This strategy, however, has not been able to avoid recent catastrophic fires (Figure 19.4). It has even been argued that it is the extinction of fires that actually causes catastrophic fires (Minnich 1983, 2001). This controversial hypothesis has been called the paradox of extinction. Its rationale is that the systematic extinction of all wildfires allows a build-up of fuel that will be consumed in future large fires caused by very adverse weather. Without fire suppression, there are frequent and small fires, while fire suppression leads to fewer, larger fires. According to this hypothesis, modern large fires arise as an artefact of fire suppression, and they can be prevented by the creation of a landscape mosaic of patches of different ages. The paradox of extinction has generated a strong debate in California in particular. Other researchers argue that the probability of having large fires would be independent of the fire suppression efforts and would respond mainly to extreme weather situations (Keeley et al. 1999; Keeley and Fotheringham 2001; Moritz et al. 2004). In most countries the relevant authorities encourage people to be careful (Figure 19.10) and not to ignite fires

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in forested areas, and they prosecute those that ignore this advice. Again, the success of this policy is open to question (Figure 19.2) as a reduction in the number of ignitions would obviously reduce the number of fires, but not necessarily the area burnt or the number of large fires in the longer term. The only way left to managers to fight wildfires is by reducing the amount of accumulated fuel, by clearing, burning, or grazing. Grazing and wood collection successfully explain the difference in the area burnt between Mediterranean southern Europe and North Africa (Figures 19.1 and 19.2). The decline of these traditional practices also explains the sudden increase in the area burnt in Mediterranean countries of Europe that occurred in the 1950s and 1960s with the abandonment of rural areas. The example of Spain in this respect is shown in Figure 19.2. However, people will not go back to the countryside, at least to work there, in the same way that their parents or grandparents did. Fuel stocks will stay in place until they are thinned or until the next fire occurs. This fire can be a prescribed fire (a low-intensity fire that is deliberately ignited by skilled teams and burns under controlled conditions), a managed fire (a wildfire that is kept burning freely as long as it is considered to be under control), or a wildfire. This problem is not exclusive to Mediterraneantype areas, but affects many other regions in the developed world. Much of the western USA, for example, has an overwhelming excess of forest fuels (Agee and Skinner 2005). In practice it is very difficult to isolate the effect of each process, namely the number of ignitions, fuel buildup, fire suppression, and meteorological conditions. An experimental approach does not help when large areas and long periods of time are involved, as is the case when dealing with fire regimes. Thus, support for the prevalence of the various component processes usually comes from indirect observations. A comparison of areas with and without active fire suppression from southern California (USA) and northern Baja, California (Mexico) seemed to provide support for the dominant role of fuel build-up (Minnich 1983; Minnich and Chou 1997). However, charcoal accumulation in sediments in Santa Barbara basin (Mensing et al. 1999), and the analysis of a Fire History Database (Keeley et al. 1999) provided support for the assumption that extreme weather events are the key drivers of large fires in Mediterranean California. However, these conclusions can be criticized for several reasons. First, charcoal accumulation may reflect a runoff cycle rather than a fire cycle (Minnich 2001). Second, fire-suppression policy is not the only difference between southern California and Baja California

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Fig. 19.10. A fire prevention sign in the uplands of Majorca near Luc. The sign is in English for the benefit of tourists (photo: Jamie Woodward).

(Keeley et al. 1999). Third, during the twentieth century other variables have changed (e.g. land fragmentation and climate) besides the increase in the suppression effort. These questions can be addressed using simple models of vegetation and fire spread over homogeneous areas. Such models incorporate meteorological variability, different rates of fuel accumulation, number of ignitions per year, fire-fighting capacity, and prescribed burning. One such model shown in Figure 19.11 was calibrated against fire regime data (number of fires and area burnt per year and distribution among size classes) of northeast Spain, central Portugal, France, and California (Piñol et al. 2005). According to this model, the total burnt area is more or less the same despite any effort to reduce it by extinguishing fires or by using prescribed burning. Nevertheless, the effect of the fire exclusion policy slightly enhances the dominance of large fires, whereas the use of prescribed fires greatly reduces the importance of large fires. In another modelling exercise a similar response was found by varying the number of ignitions—fewer ignitions only slightly reduced the area burnt, but this tends to increase the dominance of large

fires. This approach has been developed further by Piñol et al. (2007) to model fire regime characteristics in a variety of Mediterranean contexts and to assess the degree of uncertainty in the model predictions. Taking into account that large fires are, by far, the most destructive to nature, property, and human lives, it would be more efficient to increase the low rate at which prescribed burnings are conducted in Mediterranean Europe than to invest more resources into increasing the already considerable fire-fighting capacity. Prescribed fires are routinely applied in France (6,000 to 10,000 ha per year) to pastures and to forest understorey. Experimental programmes of prescribed fires are occasionally set in Italy, Portugal, and Spain. In North Africa and the eastern Mediterranean this practice is not carried out and in some countries such as Greece it is against the law to set a prescribed fire. In contrast to other continents, in many Mediterranean countries public opinion does not look favourably on these practices because of the long history of fire being perceived as a catastrophe. In fact, historically, fire has been used for multiple purposes and it is still a useful tool for some management objectives, such as the improvement of habitats and grazing areas

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(c) Fig. 19.11. An illustration of the fire regime model of Piñol et al. (2005) and of its simulation capabilities. (a) Map showing the age of vegetation after 600 years of a simulation calibrated against the fire regime data of Tarragona in north-east Spain. (b) The proportion of the age of the vegetation in the same simulation between years 300 and 600. Vertical black lines indicate the area burnt each year. (c) The proportion of the age of the vegetation between years 300 and 600 in the same simulation but with the proportion of prescribed fire increased to 1 per cent of territory per year. The lightest shading represents recent fires, light grey represents young vegetation, and dark grey represents old vegetation.

or to facilitate the regeneration of some species (Chapters 7 and 23). In this context fuel reduction is the main goal that is achieved by fires themselves. Alternatively, managed fires may also contribute to fuel reduction, as prescribed fires do.

The Future In the previous section the importance of fuel accumulation in shaping the fire regime in Mediterranean regions was stressed. However, climate has also changed during the course of the twentieth century in the Mediterranean and it is expected to continue in the same direction during the twenty-first century in accord with the dynamics of global climate change (Chapter 3). Recent studies suggest that, if the current climatic trends remain constant, fuel conditions in summer will become drier each year and the risk of large fires will increase (Piñol et al. 1998; Pausas 2004). Future climatic scenarios based on global circulation models also predict warmer summers and winters and increased potential evapotranspiration in the Mediterranean basin (Palutikof and Wigley 1996; IPCC 2001; Chapter 3). Although there is no clear agreement on future precipi-

tation trends, it is expected that there will be an increase in rainfall variability (Summer et al. 2003) and, particularly, in the frequency of extreme drought episodes (IPCC 2001). Thus, it is very likely that the climatic firerisk will increase in the near future in the Mediterranean (Mouillot et al. 2002). Whether this increased fire risk will result in a change in the fire regime of the region is another question, as there are many complex interactions between climate, land use, management policies, and vegetation as this chapter has shown. However, the increase in the wildland-urban interface and the further abandonment of marginal agricultural land all points to greater fire activity in the Mediterranean region. Depending on the magnitude of this increased fire activity, changes in the dominant vegetation types may occur, as some species will not be able to replenish their seed banks or resprout with the same vigour after successive fires. Díaz-Delgado et al. (2002) used satellite imagery to show that the recovery of vegetation (measured in terms of NDVI) slowed down after successive fire events. This effect would act as a negative feedback on fire activity, as new fuels will become available more slowly than before. In summary, the prospects are rather gloomy for the Mediterranean region in this respect. In areas where

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the vegetation is able to cope with the likely increase in meteorological fire risk, then more intense fires will be likely to occur. In areas where vegetation recovery is reduced, the expectation is a reduction in the structure and biomass of the vegetation as well as changes in the dominant species (Chapter 23). These shifts may be enhanced by changes in land use promoted by rural communities as a result of fire, as observed in central Catalonia, where burnt Pinus nigra forests do not regenerate and are replaced by grassland and deciduous forests (Espelta et al. 2002). An alternative is that a sound fire management policy would be able to allow more fires to occur at low intensity (prescribed and managed fires), instead of the destructive high-intensity wildfires waiting in the near future. Fire is a complex phenomenon and its management cannot rely on simple premises. Several strategies must coexist according to specific goals, spatial and temporal scales, and regional idiosyncrasies. Thus, fire-fighting is the first priority when lives and state properties are in danger, but it will not prevent future fires. Instead it will increase the fuel load that will eventually increase the likelihood of catastrophic fires. Fuel reduction by means of prescribed fires cannot be applied to the whole of the Mediterranean region, but it may be an efficient way to keep fuel loads low at particularly sensitive sites such as critical areas for fire propagation. Extensive fuel removal by herbivores or clearing practices may also contribute to reductions in fire extension and intensity. Avoiding non-native flammable species in forests would also help. It is clear that the simplistic view that considers fire as a disaster that can be defeated by improving fire-fighting techniques and by reducing the number of fire ignitions is not realistic. Instead the ultimate causes of the whole phenomenon are structural and due to drivers that determine the outcome of each single event, namely: climate change, vegetation succession, and social changes.

Acknowledgements This chapter is dedicated to the memory of people that lost their life fighting wildfires. We thank Marta Miralles for assistance in collecting statistical data. The research carried out by the authors and reported in this chapter was partially funded by the Spanish MCYT project REN 2003-07198, and the European Project EUFIRELAB. We thank Jamie Woodward and the external referee for reviewing our chapter. The Fire-fighting Corps of the Department d’Interior of the Generalitat de Catalunya also provided support in the development of this chapter.

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Mouillot, F., Rambal, S., and Joffre, R. (2002), Simulating climate change impacts on fire frequency and vegetation dynamics in a Mediterranean-type ecosystem. Global Change Biology 8: 423–37. Palutikof, J. P. and Wigley, T. M. L. (1996), Developing climate change scenarios for the Mediterranean region, in L. Jeftic, S. Keckes, and J. C. Pernetta (eds.), Climate Change and the Mediterranean. Arnold, London, 27–56. Pantis, J. D. and Mardiris, T. A. (1992), The effect of grazing and fire on degradation processes of mediterranean ecosystems. Israel Journal of Botany 41: 233–42. Pausas, J. G. (2004), Changes in fire and climate in the Eastern Iberian Peninsula (Mediterranean Basin). Climatic Change 63: 337–50. Carbó, E., Caturla, R. N., Gil, J. M., and Vallejo, R. (1999), Post-fire regeneration patterns in the eastern Iberian Peninsula. Acta Oecologica 20: 499–508. Pereira, J. M. C., Sa, A. C. L., Sousa, A. M. O., Silva, J. M. N., Santos, T. N., and Carreiras, J. M. (1999), Spectral characterisation and discrimination of burnt areas, in E. Chuvieco (ed.), Remote Sensing of Large Wildfires in the European Mediterranean Basin. Springer, Berlin, 123–38. Piñol, J., Terradas, J., and Lloret, F. (1998), Climate warming, wildfire hazard, and wildfire occurrence in coastal eastern Spain. Climatic Change 38: 345–37. Beven, K., and Viegas, D. X. (2005), Modelling the effect of fire-exclusion and prescribed fire on wildfire size in Mediterranean ecosystems. Ecological Modelling 183: 397–409. Castellnou, M., and Beven, K. J. (2007), Conditioning uncertainty in ecological models: assessing the impact of fire management strategies. Ecological Modelling 207: 34–44. Quilès, D., Rohr, V., Joly, K., Lhuillier, S., Ogereau, P., Martin, A., Bazile, F., and Vernet, J. L. (2002), Les feux préhistoriques holocènes en montagne sub-méditerranéenne: premiers résultats sur le Causse Méjean (Lozère, France). Comptes Rendues Palevolution 1: 59–65. Riaño, D., Chuvieco, E., Salas, J., Palacios-Orueta, A., and Bastarrika, A. (2002), Generation of fuel type maps from Landsat TM images and ancillary data in Mediterranean ecosystems. Canadian Journal of Forestry Research 32: 1301–15. Ricotta, C., Arianoutsou, M., Díaz-Delgado, R., Duguy, B., Lloret, F., Maroudi, E., Mazzoleni, S., Moreno, J. M., Rambal, S., Vallejo, R., and Vázquez, A. (2001), Self-organized criticality of wildfires ecologically revisited. Ecological Modelling 141: 307–11. Riera-Mora, S. and Esteban-Amat, A. (1994), Vegetation history and human activity during the last 6000 years on the central Catalan coast (northeastern Iberian Peninsula). Vegetation History Archaeobotany 3: 7–23. Rothermel, R. C. (1983), How to predict the spread and intensity of forest and range fires. USDA Forest Service, National Wildlife Coordinating Group, General Technical Report INT-143. Ruiz-Gallardo, J. R., Castano, S., and Calera, A. (2004), Application of remote sensing and GIS to locate priority intervention areas after wildland fires in Mediterranean systems: a case study from south-eastern Spain. International Journal of Wildland Fire 13: 241–52.

Sadori, L. and Giardini, M. (2007), Charcoal analysis, a method to study vegetation and the climate of the Holocene: The case of Lago di Pergusa (Sicily, Italy), Geobios 40: 173–80. Salvador, R., Valeriani, J., Pons, X., and Díaz-Delgado, R. (2000), A semiautomatic methodology to detect fire scars in shrubs and evergreen forests with Landsat MSS time series. International Journal of Remote Sensing 21: 655–73. Scott, D. F. (1993), The hydrological effects of fire in South African mountain catchments. Journal of Hydrology 150: 409–32. Singh, G., Kershaw, A. P., and Clark, R. (1981), Quaternary vegetation and fire history, in A. M. Gill, R. H. Groves, and I. R. Noble (eds.), Fire and the Australian Biota. Australia. Australia Academy of Sciences, Canberra, 3–21. Summer, G. N., Romero, R., Homar, V., Ramis, C., Alonso, S., and Zorita, E. (2003), An estimate of the effects of climate change on the rainfall of Mediterranean Spain by the late twenty first century. Climate Dynamics 20: 789–805. Tàbara, D., Saurí, D., and Cerdan, R. (2003), Forest fire risk management and public participation in changing socioenvironmental conditions: a case study in a Mediterranean region. Risk Analysis 23: 249–60. Terradas, J., Piñol, J., and Lloret, F. (1998), Risk factors in wildfires along the Mediterranean coast of Iberian Peninsula, in L. Trabaud (ed.), Fire Management and Landscape Ecology. International Association of Wildland Fire, Fairfield, Washington, 297–303. van Wagner, C. E. (1987), Development and Structure of the Canadian Forest Fire Weather Index System. Forestry Technical Report 35. Canadian Forestry Service, Ottawa. van Wilgen, B. W., Bond, W. J., and Richardson, B. M. (1992), Ecosystem Management, in R. Cowling (ed.), The Ecology of Fynboss. Nutrients, Fire and Diversity. Oxford University Press, Cape Town, 345–71. Vázquez, A. and Moreno, J. M. (1998), Patterns of lightning- and people-caused fires in peninsular Spain. International Journal of Wildland Fire 8: 103–15. (2001), Spatial distribution of forest fires in Sierra de Gredos (Central Spain). Forest Ecology and Management 147: 55–65. Vélez, R. (2000), Los incendios forestales en la cuenca mediterránea, in R. Vélez (coord.), La defensa contra incendios forestales. Fundamentos y experiencias. McGraw-Hill, Madrid, 3.1–3.52. Viegas, D. X. (2004), Cercados pelo fogo. Os incêndios florestais em Portugal em 2003 e os accidents mortais com eles relacionados. Minerva, Coimbra. Viegas, M. T., and Ferreira, A. D. (1992), Moisture content of fine forest fuels and fire occurrence in Central Portugal. International Journal of Wildland Fire 2: 6–86. Sol, B., Bovio, G., Nosenzo, A., and Ferreira, A. (1994), Comparative study of various methods of fire danger evaluation in Southern Europe, in Proceedings II International Conference Forest Fire Research. Coimbra, ii. C.05, 571–90. Piñol, J., Viegas, M. T., and Ogaya, R. (2001), Estimating live fine fuels moisture content using meteorologically-based indices. International Journal of Wildland Fire 10: 223–40. Viney, N. R. (1991), A review of fine fuel moisture modelling. International Journal of Wildland Fire 1: 215–34.

This chapter should be cited as follows Lloret, F., Piñol, J., and, Castellnou, M. (2009), Wildfires, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 541–558.

IV

Environmental Issues in the 21st Century

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Editorial Introduction Jamie Woodward

This volume has traced the development of the Mediterranean landscape over very long timescales and has examined modern processes in a wide range of settings. Earlier chapters have explored tectonic processes and the evolution of the topography and biota, the nature and impact of Quaternary climate change, and natural hazards, as well as the increasing role of human activity in shaping geomorphological processes and ecosystems during the course of the postglacial period. A core theme in several chapters is the nature of the relationship between humans and the Mediterranean environment. Over the last one hundred years or so, and especially in the period since the Second World War, this relationship has changed dramatically. Resource exploitation, urban expansion, and rural depopulation have all taken place at unprecedented rates, with major impacts upon the quality of land, water, air, and ecosystems. The final part of this volume examines four key topics of environmental concern; its four chapters explore, respectively, land degradation, water resources, interactions between air quality and the climate system, and biodiversity and conservation. Where possible, it is important to place these issues within an appropriate historical perspective. Many components of the Mediterranean environment have responded in a sensitive way to past environmental changes, but the pressures on land and water resources have never been more intense. Improved monitoring networks and new modelling efforts are needed to predict more effectively the impact of climate and social change on all environmental systems and to help inform policymakers seeking a more sustainable use of the region’s resources. Chapter 20 examines the ecological aspects of land degradation and sets out new ideas on productivity dynamics. It explores some of the interactions between land use change, vegetation dynamics, grazing patterns,

and wildfires. The uneven geography of water resources and water use are highlighted in Chapter 21. Water resource issues have become an increasingly important factor in the geopolitics of the region against a background of climate change uncertainty, rising demand, and a diminishing resource base. Chapter 22 analyses the interactions between climate, air quality, and the water cycle. Summertime air pollution adversely affects human health throughout the basin; the stable and stratified summer atmosphere and intense solar radiation—in combination with local pollution circulation systems—is described as a ‘cooking vessel of photochemical smog’. Ozone production, aerosol haze, and pollution plumes from Asia and central Europe are major air quality issues. The final chapter in this book examines biodiversity and conservation across the region—it traces the development of ideas and draws together research from a range of terrestrial and aquatic habitats. Biodiversity and habitat quality are closely related to land, water, and air quality. A key feature of Mediterranean ecosystems is the deep history of interaction with humans— especially through the early use of fire as a management tool and the fundamental changes that accompanied domestication and the development of agriculture. In more recent times, unchecked urban development, especially along the coastal fringe, poses a major challenge to the conservation agenda. Distinctive hotspots of biodiversity with many rare and endemic species have been identified within the basin. These are partly a product of high topographic variability and isolation and illustrate very clearly the fundamental and long-standing links between geodiversity and biodiversity in the Mediterranean world. The protection of these landscapes and habitats is a key challenge for the twenty-first century.

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20

Land Degradation John Thornes

Introduction ‘Land degradation’ means the reduction and loss of the biological or economic productivity caused by land use change or by a physical process or a combination of the two. ‘Land’ means the terrestrial bio-productive system that comprises soil, vegetation, and other biota and the ecological and hydrological processes that operate within the system (UNEP 1992). The main components of land degradation are ecological degradation, soil loss, and reduction in the amount and quality of the available water resources for human survival and economic sustainability. Conacher and Sala (1998) have edited a major volume on land degradation in Mediterranean environments of the world and soil erosion mechanisms and water resources are considered in other chapters of this book (Chapters 6 and 21). This chapter will focus on the ecological aspects of land degradation by exploring some of the interactions between land use change, vegetation dynamics, grazing patterns, and wildfires. This chapter will also try to identify and avoid repeating the myths that abound in the more popular and/or politically motivated accounts of Mediterranean land degradation.

A Brief History of Land Degradation in the Mediterranean Because of the complex spatial mosaic of environmental and cultural conditions across the Mediterranean (see Blondel 2006), it is not simple to identify the causes or main controls of land degradation or the management strategies required to combat degradation (Lesschen et al. 2007; Märker et al. 2008). As discussed in the context of lake sediment records in Chapter 9,

it is certain that the origins of land degradation extend far back into prehistory. Indeed, Naveh and Dan (1973) have proposed a seven-phase history of land degradation for the Mediterranean basin, paraphrased thus: Phase 1 was the Lower Palaeolithic (around 1,000,000 to 100,000 years BP), when the Levant was the main route of biotic and hominid dispersal from Africa to Eurasia and later westwards through the Mediterranean basin. Hunting and gathering were the main activities and the populations were probably very low. Human impact on the environment is not known—but land degradation is assumed to have been negligible. After this, in Phase 2, it is argued that the use of fire as a tool for the opening up of dense forest spread westwards from Greece, possibly reaching France as early as 400,000 BP. Most authors agree that the Neolithic period (Phase 3) was a major time of landscape change. This was the period of domestication of plants, animals, and domestic livestock (Blondel 2006), that again progressed from east to west, being 4,000 years earlier in the eastern than in the western part of the basin (Chapters 5 and 9). During this period, human population increased more than tenfold—a rate of increase not seen again until the twentieth century. Around 4,000 BP, the climate became more arid in southern France and the breeding of sheep and goats was definitely established (Le Houérou 1981). Moreover, transhumance between the garrigue of Languedoc and the upland pasture of the Cévennes was already in progress. Phase 4, from 5,000 BP to the end of the Roman period, saw the domestication of fruit trees, land clearances, and extensive terracing on hillslopes (Chapter 7). Soil and water conservation methods were implemented, for example in Tunisia and in Israel (Lavee et al. 1997). Populations increased and the cutting down of trees for shipbuilding began. In Phase

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5 the Moslem conquest was, according to Naveh and Dan, accompanied by economic decline in the Levant as the pastoral nomadism of the Arab tribes replaced the developed hillsides and irrigation systems. By contrast, in the Moorish kingdoms of Spain, irrigation was more thoroughly planned and land conservation notably improved by the Arab invaders. Phase 6 of Naveh and Dan is the technological phase from the late nineteenth century, with drainage of swamps, land reclamation, monoculture in agriculture and forest, and the development of large-scale irrigation schemes across the Mediterranean basin. This culminates in a Phase 7 in the twentieth century with the mechanization of agriculture and its industrialization. Le Houérou (1981) disputes the concept that the pastoral and Bedouin civilizations had a greater effect on vegetation and land degradation and believes that the Hellenistic–Roman period had a more drastic impact due to the large populations. Between classical and recent historical periods, Le Houérou argues for a relative stabilization of Mediterranean vegetation, thanks to the slowing down of demographic growth (world population barely doubled from AD 640 to 1650). In Palestine and North Africa, he argues, population actually fell. According to Despois (1961), it seems certain that North Africa has never, in the course of its history, seen as great an extension of crop lands, nor such an increase in population as at the present time. Le Houérou (1981: 481–2) believes that ‘[i]ts vegetation and soils have never been at such a risk’ and that ‘[e]rosion and desertification progress at a terrifying rate. The degradation is essentially the result of human activity; climate does no more than provide favourable, though constant conditions.’ Of course there is no reason to expect a uniform history of degradation across the entire basin. As our knowledge improves in local areas, more complicated histories can be expected to emerge. An excellent example is the study of the Marmora plateau to the northeast of Rabat, Morocco, by Nafaa and Watfeh (2002). They describe the past environmental changes in what is known as the Marmora Forest. The plateau is sprinkled with a forest of cork oak, which used to be one of the largest forests in Morocco, but is gradually being eliminated. Today 180,000 sheep and 6,000 cattle spend more than 300 days per year in the forest, which is considered to be overgrazed. Before 1913, 100,000 trees had been felled. But the degradation of the Marmora Forest cannot entirely be attributed to contemporary exploitation. Studies of the stratigraphy at the coast have shown that, in the Soltanian period (c.22,000– 26,000 BP), the sands overlying the plateau were

reworked extensively by wind action under drought conditions. This is taken to imply that the forest was not in existence at that time, but it appears to have regenerated into a cork oak forest after about 20,000 BP. Moreover, during the Holocene, there is stratigraphic evidence of an alternating sequence of aeolian and colluvial conditions on the slopes, again implying an alternation of steppe and forested conditions (Chapters 4 and 9). As local ecological histories become better known, it seems likely that a series of changes from stable forested conditions to unstable wind-eroded steppe conditions will be revealed. In the last ten years or so, Mediterranean land degradation has come to the political stage for the following reasons: r It has become increasingly obvious that global

r r

r

r

warming, accompanied by regional effects, has become a reality rather than speculation (Chapter 3). The flight from rural areas and the downward spiral of rural services continues unabated after centuries. A series of natural disasters—floods and the drought of the 1980s—have highlighted perceptions of the sensitivities of rural and urban areas to environmental change. The European Union Agenda 2000, seeking to rectify problems with the Common Agricultural Policy, has identified land degradation as a major problem and this has been coupled to a review of soil as a major renewable resource that is impeding the development of a policy for sustainable livelihoods. The EU Water Framework Directive (WFD) (2000) seeks to provide tools for the management and control of river systems using the principles of integrated catchment management. This requires that all river basins are provided with management plans for the monitoring and control of flooding, water quality, and consumable water resources—including groundwater (Chapters 8, 18, and 21). Land degradation within a river basin is important for all of these and therefore necessarily must be part of the monitoring and planning under the Water Framework Directive. Thornes and Rowntree (2006) have expressed some misgivings about the suitability, applicability, and desirability of the unselective application of the WFD across the European Mediterranean countries mainly on the grounds that it is designed for the temperate environments of north-west Europe.

Land Degradation

Contemporary Perceptions of Land Degradation Land degradation in the Mediterranean has been a serious topic since classical times and has been the subject of legislation in all the European Mediterranean countries and in the Maghreb (Burke and Thornes 1998). One example is from Italy, where the most important law for environmental resources, dating from 1989, deals with the reorganization and operation of soil conservation (Freschi 1998). The law defines the planning of activity in a river catchment by means of a basin plan and new central and peripheral organizations are charged with planning and carrying out operations to provide for rational use of water resources and to safeguard buildings and infrastructure from landslides, avalanches, and erosion. A second example is from Morocco, which ratified the United Nations Convention to Combat Desertification (UNCCD) in 1996. Specific measures have been taken to prevent and/or reduce land degradation, to rehabilitate partly degraded land, and to reclaim desertified land (Price 1998). In his report, Price gives a comprehensive account of the threats of, and policy against, land degradation in Morocco, indicating the high level of national awareness of the problem. If further research into local ecological histories reveals a series of shifts from bare ground to a healthy

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vegetation cover, then a single cycle of the r–K sequence of earlier ecologists might provide a suitable model for the Mediterranean. This approach envisages a change from a clear unvegetated state occupied by exploratory, colonizing plants (r-species) through to a stable cover of mature trees occupied by conservative, more woody plants that are less easily perturbed (K-species). In a recent development of this basic model, Gunderson and Holling (2002) envisage two further phases of collapse (the Ÿ phase) and reorganization (the · phase). In developing this scheme, called ‘panarchy’, they hypothesize the existence of multiple adaptive cycles in time and at different spatial scales as shown in Figure 20.1. The trajectory from r to K replaces the conventional hypothesis of succession and is the logistic curve of modern ecology. The panarchy hypothesis seems to offer a more robust and flexible approach to explaining the history of degradation in the Mediterranean region and has the advantage that its intrinsic dynamics do not need to call upon ubiquitous explanations for growth, collapse, and regeneration of environmental systems that were outlined above. It also avoids the general unified theory approach where a single explanation for degradation is sought for the whole of the Mediterranean region. Because of the increased availability of satellite data gathering and other achievements of modern

Fig. 20.1. Different phases of the Holling and Gunderson adaptive cycle where r is the colonization phase, K is the consolidation phase, Ÿ is the release phase and · is the system reorganization phase. See text for discussion. Modified from Holling and Gunderson (2002).

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technology, the new data age has given rise to widespread concern over land degradation. This has coincided with the centralization of European bureaucracy and, based on the assertion that land and livelihoods are faced by a severe crisis, has led to a call for a battle against degradation around the Mediterranean. What one might call the ‘Grand European Desertification Project’ started in Mytilene in the Aegean in 1984, with a conference on European Desertification (Fantechi and Margaris 1986). This grew through a series of massive projects in successive research frameworks and is now moving towards a major programme for the mitigation of land degradation across Europe. This is based on the continual assertion of national governments from Greece to Portugal that land degradation is severely undermining planning to stabilize rural environments (Agenda 2000). There is a firmly held belief that the southern countries of Europe are gripped in a crisis of land degradation that can only be exacerbated by the coming effects of global warming (IPCC 2001). Following the UNCOD Nairobi Conference on Desertification in 1977, recognition of the problem in North Africa and the Near East led to ratification of the United Nations Convention on Combating Desertification (UNCCD) and adoption of the associated actions. The problem, as Le Houérou (1981: 482) had previously expressed it in relation to the Maghrebian states, is that ‘Erosion and degradation progress at a terrifying rate.’ There are arguments both against and for this view. In recent years, land degradation has been encompassed by the term ‘desertification’ and thereby has taken on all the uncertainty and political overtones that this term implies (Reynolds and Stafford Smith 2002). Land degradation is a more satisfactory term that avoids many of these ambiguities and difficulties, especially when applied to the Mediterranean. Though ‘desertification’ cannot be regarded as ‘mythical’ (cf. Thomas and Middleton 1994), statements about the intensity and extent of land degradation are at best inaccurate and at worst based on nothing better than guesswork. This is because land degradation in the Mediterranean has been evaluated mainly by soil erosion, which is virtually impossible to measure with any degree of confidence. It is normally expressed as tonnes per hectare and measured on field plots as runoff and sediment yield. The plots are often inadequately designed, poorly located, and operated over too small a time span to produce meaningful results. They are, after all, only a sample in time and space. For this reason, the larger the sample size and the greater the range of conditions covered, the more likely it is that the results will be useful for practical

purposes. So while they often provide valuable experimental results, they hardly provide a basis for policy decisions. One issue in this respect is that it is the loss of productivity that is critical. In the early stages of land degradation—such as the period following land abandonment—because soil nutrients are concentrated in the upper horizons of the soil, organic matter and nutrients are lost in considerable amounts. Although the loss of soil may be small, the loss of productivity can be quite large and land degradation in the wider sense becomes critical. In the search for simple one-number indicators, the amount of soil lost is often advocated as the most valuable indicator of land degradation when that is probably not the case. Perversely, low rates of erosion, when discovered by researchers, may be dismissed as inaccurate because of expectations that Mediterranean rates are much higher than those in temperate environments and the perception of the Mediterranean as a ruined landscape or lost Eden. The reasons for, and history of, this misconception are discussed in Grove and Rackham’s The Nature of Mediterranean Europe: an Ecological History (2001). As well as data from global modelling, empirical data do, in fact, reveal that the Mediterranean is a region of relatively high soil erosion (Chapter 6), even recognizing the very high spatial variability. So far, most estimates of erosion have been based on the USLE (Figure 20.2). The underlying approach has also been used for the estimation of soil losses based on the CORINE (Coordination of Information on the Environment) GIS system. The maps produced and the accompanying table indicate the actual erosion risk in the European Mediterranean (Table 20.1). This seems to indicate that the largest absolute area of high-risk land is in Spain, with about 30 per cent of land in the high risk area (Figure 20.3). The actual soil erosion risk map shows that the risk is TABLE 20.1. National soil erosion risk data for five Mediterranean countries in the EU Country

High risk area km2

%

Moderate risk area km2

France (south) 1,693 1 22,362 Italy 30,169 10 93,983 Greece 27,713 19 47,877 Spain 145,494 29 219,908 Portugal 26,878 30 48,166 EC (south) 228,947 19 432,295

Low risk area

%

km2

%

12 31 36 44 54 36

123,643 165,823 39,287 111,518 12,884 453,155

65 55 30 23 15 37

Excluded area km2

%

42,463 22 11,303 4 20,113 15 20,598 4 1,000 1 95,477 8

Note: The data for Spain and Portugal do not include the islands. The excluded area column comprises urban land, lakes, bare rock terrains, and areas without data.

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567

Fig. 20.2. An application of the USLE for the Autonomous Region of Andalucia, southern Spain. Source: Ministerio de Media Ambiente, Junta de Andalucia. Note the very high rates of erosion in the east of the area near Almeria (Chapter 1).

concentrated in Portugal and southern Spain. All these data must be viewed with caution, though they are the best available estimates.

Land Degradation: The Main Processes The Role of Vegetation Land degradation, as defined above, is absolutely and inextricably linked to the vegetation cover and functional type. The vegetation cover is the percentage of the ground that is covered by vegetation when viewed vertically from above. This directly controls the amount of erosion that occurs in relation to the bare ground value (Figure 20.4; Chapter 6). As the percentage of cover decreases, the amount of erosion increases. This increase in erosion is most rapid just below 30 per cent vegetation cover. Salvucci and Eagleson (1992) have demonstrated that the ecologically optimal cover

in dry lands with this value occurs when the ratio of actual to potential evaporation is also about 30 per cent. Below this, erosion rates increase catastrophically (Elwell and Stocking 1976; Francis and Thornes 1990). Functional type refers to whether the plants are grasses, shrubs, or trees and the functional type depends on the trade-offs that the plants have to make in capturing and storing different resources, notably light and water (Tilman 1982). This means that where water is scarce and light is abundant, more carbon is allocated to roots and stems than to leaves. Mulligan (1998) has shown that these functional types affect erosion rates. Human exploitation of the vegetation cover also depends on the functional types—grazing animals have a strong preference for leaves, firewood gatherers for woody stems and branches. The vegetation evolves in response to the controls of light, water, and temperature to reach a steady state in pattern, amount, and functional type. Erosion responds accordingly. In a stable ecosystem, physical or

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Percentage of land at high erosion risk 80 – 100 60 – 80 40 – 60 20 – 40 0 – 20

Fig. 20.3. The EU erosion estimate for Spain based on the CORINE database showing the percentage of land at different levels of erosion risk. See text for discussion.

human perturbations have little effect and the system can restore itself to the previous steady state. A simple metaphor (Figure 20.5) is to regard the values of the system state variables as a landscape over which a ball is rolling. The ball comes to rest in the depressions and rolls away from the bumps. The depressions (attractors) are those values of the system variables where stability prevails (the ball stops moving). By contrast, the bumps are unstable locations that the ball moves away from (repellers). If the ball is in a depression energy is needed to move it to another position (changing its state). The deeper the depression, the more energy is needed to change the state and the more stable is the system. On a flat plain the ball rolls about at random (varies randomly in time, but only by tiny amounts). A fuller

description of this metaphor is given in Holling et al. (2002a ). Unstable states give large fluctuations through time in the vegetation cover (and hence in erosion). Generally speaking, droughts provide instability in Mediterranean ecosystems as the vegetation tries to reach ecological optimality with respect to the changing hydrological conditions. Under severe perturbations, the system may start to oscillate widely with large underand over-adjustments. This situation was modelled for dry Mediterranean conditions by Thornes and Brandt (1994) using real rainfall records and a boom-andbust cycle in vegetation cover and erosion was the result. Pueyo et al. (2007) have argued that land use fragmentation and aridity were the key controls on

100

100

75

75

Soil loss

50

50

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25 Runoff 0

0 0

25 50 75 Mean seasonal vegetation cover (%)

Runoff as a percentage of mean annual rainfall

Soil loss as a percentage of bare plot results

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(a)

(b)

100

Fig. 20.4. Soil loss and runoff as a function of the proportion of ground covered by a vegetation canopy (after Elwell and Stocking 1976).

the extent of vegetation recovery between 1957 and 1998 in parts of the middle Ebro valley in north-east Spain. Interannual rainfall variability is greater in dry Mediterranean regions (Chapter 3), therefore the ecosystems there appear to be more unstable. As the climate becomes drier or wetter the interannual variability reduces, as does the magnitude of erosion. After fire in the more humid Mediterranean environments, such as the Catalan countries of Spain and the mountain areas of Greece and Italy, field experiments appear to indicate that after 8–10 years the rate of sediment production typically falls back to levels that existed before the fire. In drier areas, with greater vegetation instability, the impacts of fire should be greater and longer-lasting, especially if the vegetation cover is close to the 30 per cent cover level. (See Chapters 6 and 7.) The overall conclusion is that, where wash erosion dominates, rates should be highest where it is climatically drier, but not arid. This reinforces the opinion that rates of erosion are highest in dry Mediterranean areas with rainfall around 250 mm and with actual evapotranspiration less than 30 per cent of potential evaporation. This matches with Langbein’s and Schumm’s (1958) empirical curve of sediment yield for the conterminous United States that shows a peak of sediment yield at approximately 300 mm per year (Figure 20.6a). Kirkby’s (1980) model-based analysis shows a similar peak for wash erosion (Figure 20.6b). On the semi-arid to arid side of this curve, as runoff decreases, so too does sediment yield. On the wetter side there is a very

(c)

(d)

a b

Metaphor

Phase space

Fig. 20.5. A metaphor for a system’s stable and unstable conditions (modified after Gunderson and Holling 2002); (a) is a level surface, (b) has a slight depression, (c) has a hill, and (d) has two hills and two depressions. See text for discussion.

steep fall to lower suspended sediment yields as the vegetation cover increases and shifts from shrubland to grasses and eventually to forest. Kosmas et al. (2002) have gathered a large sample of erosion plot data that appears to confirm the general behaviour described

570

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(a) 20

1,000 Forest 750

Grassland

500

250 Desert shrub

(c)

15

10

5

0

0

100–200 200–300 300–400 400–500 500–600

re

300 100 200 0 Annual sediment yield (tonnes/km2)

1,000

Ba

Annual erosion from a 10-m long 10° slope (mm)

Mean annual sediment yield (g m-2)

Effective precipitation (mm)

1,250

Annual precipitation (mm)

(b)

rub

Sc

100

Gras

Curves for constant vegetation cover

d

slan

Composite curve for natural vegetation

10

Bare

1

t Fores

0.1

Curves for constant vegetation cover Composite curve for natural vegetation

Water erosion

Wind erosion (very approximate)

Scrub Forest

Grassland

0.01 0

500

1,000

1,500

2,000

Mean annual precipitation (mm)

Fig. 20.6. (a) The relationship between sediment yield and annual effective rainfall (that which produces runoff) from Langbein and Schumm (1958). (b) Estimated rates of erosion by wind and water as a function of rainfall and vegetation cover (Kirkby 1980). (c) Erosion on field plots around the Mediterranean, as a function of rainfall (after Kosmas et al. 2002).

above that, at the global scale, the Mediterranean is a region of relatively high rates of erosion (Figure 20.6c) (Chapter 8).

Sediment Yield Data Given the difficulties of field-plot data referred to above, it is instructive to take note of measured suspended sediment yield data for Mediterranean rivers gathered and interpreted by Woodward (1995) who starts with a quotation from Zachar (1982: 389), ‘The greater the effects of human intervention and the more extreme the conditions of climate, the more serious the consequences of erosion have been. Thus the coun-

tries around the Mediterranean are most affected by erosion.’ The data in Woodward’s analysis appear to contradict the empirical (Langbein and Schumm 1958), theoretical (Kirkby 1980), and simulation (Thornes and Brandt 1994) results described above, but draw attention to other causes of high sediment yield. These are deep gully and badland development, exceptionally erodible lithologies, steep terrain, high-intensity rainfall and flood-prone catchments, and overgrazing by sheep and goats. He also points out that there are technical problems with the use of suspended sediment yield regressed against rainfall (spurious correlation resulting from co-variation) and, following Zachar (1982),

Land Degradation

the bias introduced by locating measurements at sites designed to measure soil loss for agricultural planning. There is, in addition, a bias towards small research basins with their very high sediment delivery ratios. The same difficulties seem to apply to reservoir siltation studies. The developments in Caesium 137 (137 Cs) analysis and stable isotope studies may provide the solution in the search for erosion estimates in Mediterranean environments (Navas and Machin 1991; Schaller et al. 2001; Zorzou 2004). In conclusion it seems that erosion rates are highly variable across space and over time and that, as yet, no reliable widely available method has been produced that enables us to compare the relative degrees of erosion in different countries across the Mediterranean basin. The most we can hope for from direct measurement of the types described is the possibility to confirm or deny the wide range of existing modelling attempts.

Wind Erosion and Salinization Two other often-emphasized degradation processes are wind erosion and salinization of soils. Wind erosion and sand movement are generally minor concerns in most natural frameworks in the European Mediterranean (Warren and Barring 2002; Riksen et al. 2003), except in small areas such as central Aragon (Gomes et al. 2003), coastal Turkey, Sicily, and Sardinia, but they are an important form of land degradation on the Saharan edge of the Maghrebian states, such as Tunisia and Algeria. In Tunisia the main area of aeolian deflation and accumulation in nebkhas and dunes is in the south-west of the country where major conservation efforts are continually carried out (Mtimet 1999). Further discussion of aeolian processes in the Mediterranean region is given in Chapter 14. Salinization is a major problem in Tunisia where saline phreatic water has destroyed 16,000 hectares in the lower Mejerda valley and the oases of Tozaur, Gafsa, Kebili, and Gabes have been significantly affected. Salt intrusion to coastal groundwater occurs where sea water invades aquifers near or at the coast (Chapter 21). The alluvial plain of Capoterra in southern Sardinia has been studied by Barrocou (1994). Salt water intrusion has resulted from over-exploitation of groundwater to satisfy the demands of agriculture, industry, and domestic use. Groundwater is withdrawn from the aquifer system from about 300 wells scattered over an area of nearly 60 km2 . Continuous monitoring has indicated a rise in salinity, particularly in the central part of the alluvial plain.

571

Groundwater Abstraction A related theme across the Mediterranean, both north and south, is the withdrawal of groundwater, notably for irrigation (Chapter 21). In Castilla la Mancha, for example, the area under irrigation increased from 15,916 ha in 1970 to 352,452 ha in 1996. There have also been reduced inputs to the aquifers through declining rainfall, especially in the general decline of the 1960s and 1970s and the dry years of the early 1980s. On top of this, a general rise in temperature has led to increasing evapotranspiration losses (Burke 1998). These hydrological budget changes have led to land degradation. In the wetlands of Tablas de Damiel (Llamas 1988) at the confluence of the Ciguela and Guadiana rivers, 20 km2 of wetland has been reduced to less than 1 km2 as a result of abstractions from the in-flowing rivers and extraction of water from the underlying aquifer. So desiccated have the wetlands become that the residual peat is also disappearing as a result of internal combustion.

Precursors of Degradation In the light of the discussion of vegetation and erosion above, it is axiomatic that changes to the canopy can be expected to lead to higher rates of erosion (Chapters 6 and 8). This is why Mediterranean land degradation is most commonly ascribed to the processes of deforestation, overgrazing, and land abandonment. The converse of deforestation as a cause of land degradation is to propose afforestation as a solution. In fact both deforestation and grazing are very variable and complicated processes. Before and after effects are very difficult to establish and paired catchments are almost impossible to compare. As Chapter 6 has shown, the underlying terrain that is exposed is also very variable as regards soil type, depth of soil, and roughness conditions.

Deforestation There is much more to the story of deforestation than is usually supposed. The meaning of ‘forest’ in historical texts is enormously variable, ranging from Mediterranean scrubland to previously planted trees. The firmest evidence of conditions before deforestation comes from long sequences of pollen records (Chapters 4 and 9). Some of the best sites for pollen preservation are in upland lakes and bog sites that, for climatic reasons, were almost invariably clothed with mature forests of woody trees with dense canopies. In some lowland areas preservation is more difficult and cave sites may poorly reflect the surrounding pollen rain. A few

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exceptional sites, such as Padul near Granada, Spain, Lago di Monticcio in southern Italy, and Lake Ioannina in north-west Greece, provide very long sequences that have been extrapolated over wide areas (Chapter 4). There are other problems in interpreting prehistoric vegetation from pollen. It can be argued that marine sediments are difficult to interpret because the origin of the pollen grains is often difficult to establish. Also, as Grove and Rackham (2001) note, palynology depends very largely on wind-pollinated plants whereas, in the Mediterranean, whole categories of plants are insect pollinated and leave very little to provide a pollen record. For an extended discussion, the reader is referred to Wainwright and Thornes (2003). Grove and Rackham (2001: 157) summarize the position for Mediterranean Europe thus: By 7000 BC there were trees in all parts of the Mediterranean that have a pollen record. Usually oaks predominated, but in Corsica the predominant tree was tree-heather and in parts of Spain, south Greece and Crete it was pine. Often there was more, sometimes much more, deciduous oak than evergreen. There is general evidence for other typical Mediterranean trees and shrubs, Pistacia (lentisic or terebinth), Phillyrea, Arbutus, plane, olive. There were also northern trees: alder, birch, elm, hazel, hornbeams (carpinus and ostrya), lime (particularly significant because it is insect pollinated).

It can be argued that, given large variations in time and space, it is extremely difficult, on the basis of the limited sample of pollen sites in what are often ecologically unique situations, to make conclusive generalizations about the state of prehistoric and early vegetation cover. However, well-dated long-term pollen-based records of vegetation change from across the Mediterranean are presented and discussed in detail in Chapters 4 and 9. There is also the problem that there have been other causes of forest loss in addition to deforestation by clearfelling. Many other forms of ecological disturbance result in the loss of woodland, such as wildfires and changes in the fallow cycle and the introduction of new animal and plant species. On some of the Greek islands, practices related to grazing and burning have also adversely affected woodland conservation. The woodland cover has been subjected to climate change that has resulted in changes in the ecologically optimal cover. There is also a problem of circularity if erosional and sedimentological changes are used to infer vegetation and climate changes in one area and the climate change is then used to explain the vegetation in another. Even if the picture of deforestation were clear, there would be some reservations about the relationship between forest clearance and degradation. Many studies

of the impact of fire reveal dramatic changes in runoff and sediment yield in the period immediately following the burn. Usually, however, after 8 to 10 years (depending on the weather sequence), the runoff and sediment yield rates have fallen back to the pre-burn levels. Another factor affecting the impact of fire and clearance is the extent to which the organic litter layer is destroyed. Brandt (1989) also showed that, for both temperate and tropical forests, the mean raindrop energy beneath the canopy was higher than that beneath open skies, presumably due to the concentration of water on leaves. Without a protective litter layer, this can yield higher soil loss rates beneath forest. Again the recovery rate and pattern are relevant (Obando 2002). In conclusion, clear felling of trees almost invariably produces increased runoff volumes and enhanced sediment yields. Nevertheless, field experiments on the impacts have proved difficult to evaluate even though they have been conducted for almost 100 years since the celebrated Wagon Wheel Gap paired catchment studies that started in southern Colorado in 1910. Prior to this experiment, it was generally assumed that: forests reduce the magnitude of ordinary seasonal floods; forests maintain stream flow in dry weather; forests prevent erosion. The essential reasons for uncertainty about the experimental outcomes—which, in general, did not support these assumptions—are discussed by Rodda (1976).

Land Abandonment Land abandonment has become a major feature of the European Mediterranean countries over the last century (Lesschen et al. 2007; Chapters 7 and 23). This has been largely attributed to rural depopulation that affected most of Europe at the same time, a mixture of declining birth rates, the failure of rural service provision, and the growth of economic opportunities. In some areas, special circumstances led to rural abandonment. Thus, in the eastern Alpujarras of Andalucia the destruction of the vine stocks by the disease phylloxera at the end of the nineteenth century had an impact on rural populations that the area never recovered from. Later, in the twentieth century, economic conditions led to further out-migration, with migrants seeking employment in large Spanish cities (notably Barcelona), in coastal resorts, and in Germany under the Gast Arbeiten scheme. In the 1980s a major drought also pushed rural population to the brink and led to a further exodus. Much of European Agricultural Policy in the last forty years has been directed to slowing or halting this flight through financial incentives (Thornes 2002).

Land Degradation 50

40

20

10

0

–10 50

Age of abandonment (years) Fig. 20.7. The evolution of plant cover over time on abandoned plots in the Spanish Pyrenees (after Garcia Ruiz et al. 1994).

There have been several studies of the relationship between land abandonment and land degradation. Garcia Ruiz et al. (1994: 28) described farmland abandonment as ‘probably the most important geo-ecological process affecting soil and water in mountain areas of Central and Western Europe’. They concluded that farmland abandonment produced many changes in runoff and sediment yield. Land management after abandonment was the most important factor in preventing any undesirable effects from a geomorphological and hydrological point of view. An important issue was the recovery of the vegetation after abandonment, dependent on the interaction between

25

Runoff coefficient (%)

20

land uses after abandonment (including grazing and fire) and the physico-chemical characteristics of the soils. Figure 20.7 shows their data on the evolution of the bush cover following abandonment. The temporal pattern shows a steady but steepening increase of bushes up to twenty-five to fifty years after abandonment, followed by a levelling off at about 30 per cent plant cover after fifty years. Figure 20.8 shows the results they obtained from plots after a period of eighteen months of observation of runoff and sediment yields. It shows very clearly that shrub cover can be an excellent plant formation for soil protection. The results also support the use of the logistic curve for vegetation recovery after abandonment, as do similar studies after fire (Godron et al. 1981). In the logistic curve there is a slow recovery at the start of recolonization (the r-phase referred to above). With feedback the process gains momentum leading to a steep rise. As resources become scarcer (water, light, and nutrients) a limit to growth sets in and the rate of increase of biomass levels off as the vegetation approaches the potential when the rate of growth falls to zero or varies around the potential. This is the K-phase of the panarchy paradigm (Figure 20.1). Field surveys in the semi-arid Caraco basin in Murcia showed that abandoned fields had more gully erosion than cultivated ones (Lesschen et al. 2007). By simulating patterns of land abandonment with four different land use scenarios for the period 2004–15, these researchers showed that, for the various land use scenarios, the potentially vulnerable areas for gully development increased from 18 to 176 hectares. Caution is needed, however, with respect to studies that show only loss of soil mass. From an agricultural point of view, what matters more is the loss of nutrients, which is a sensitive indicator of land degradation.

1,000

Runoff coefficient Sediment yield

800

15

600

10

400

5

200

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0 Shrub cover Shrub cover 100% 85%

Meadows 85%

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Plant cover (%)

30

573

Fig. 20.8. Runoff and sediment yield data for different land uses on regenerated abandoned fields in the Spanish Pyrenees (Garcia-Ruiz et al. 1994).

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A recent study by Pardini et al. (2004) carried out in the Cap de Creus Peninsula of north-east Spain, emphasized this aspect of land degradation. Small losses of topsoil can result in rapid and large losses in productivity. This fact is not often recognized in the soil erosion league tables for the Mediterranean. As Pardini and co-workers have shown, the most significant outcome is not the absolute loss of soil, but depletion of the organic carbon and total nitrogen stores—especially in burnt environments. They go on to conclude that abandoned soils left under natural vegetation are comparatively less susceptible to erosion and nutrient depletion than reforested or cultivated soils. Wildfires determine in part the sequence of shrubs that occurs (Chapter 7). When there is replacement of ericaceous shrub by Cistus shrub, this tends to create less favourable soil physico-chemical conditions and therefore increases the likelihood of degradation. In this example, forests planted with Pinus halepensis were so detrimental to soil properties that reforested areas showed lower organic content and cation exchange capacities and therefore higher susceptibilities to erosion. The regeneration after clearance by felling or following cultivation is largely affected by the weather sequence and, since agriculturalists do not have much control over the timing of their operations, regeneration is a very hit-and-miss process. Because of the global influence of ENSO (El Niño Southern Oscillation) and the connection between the North Atlantic Oscillation and the Mediterranean (Goodess and Palutikof 2002: Chapter 3), The cyclicity of Mediterranean rainfall has become more predictable. Regeneration planning could respond to these oscillations by taking advantage of runs of relatively wetter years rather than a dry series of years for planting, as observed by Holmgren and Scheffer (2001). The quicker that a vegetation cover greater than 30 per cent can be established, the less will be the resulting land degradation. Controlled experiments are difficult to achieve, but it is often helpful to carry out computer simulations in order to experiment with forest strategies. For example, in an investigation by Obando (2002), based on field studies near Lorca in south-east Spain, simulations on forest regeneration, following land clearance in a dry Mediterranean environment, examined the effects of patterns of abandonment (leading to regeneration) on the sequential characteristics of the recovery. She sought to explore the effects of four variables: different initial conditions; different patterns of abandonment; different rainfall regimes in the regeneration period; and different time-patterns of abandonment (the percentage

of land abandoned per year). She modelled runoff and erosion from a cellular grid overlying a basin topography with soil-covered slopes. The model comprised a coupling of hillslope hydrology, vegetation growth, erosion, and runoff to the outlet of the basin. The model also accommodated spatial patterning selected to allow sections of the catchment that were considered vulnerable to high erosion rates to be abandoned experimentally (e.g. areas with steep gradients or lithological crusts susceptible to erosion). Obando’s experiments produced a sequence of annual vegetation covers in the basin and the corresponding runoffs and sediment yields under the various conditions. The annual runoff generated by the model for the bare catchment is higher than that for the shrubcovered catchment over a period of twenty years. Under dry conditions (R < 100 mm) annual runoff is proportionally lower in both cases. The difference between the two catchments decreases for events >110 mm and the runoff coefficients obtained using the wet series for the bare and shrub-covered catchments are 33.5 and 21.8 per cent respectively. The daily biomass production (g/m2 ) closely tracks the rainfall and, for the bare areas (with an initial seed-bank of 0.1 g/m2 ) quickly (six years) reaches the production rate of the heavily vegetated (330 g/m2 ) plots. Because these are real rainfall series, the rate falls again after about fourteen years. In other words, in this model result, the effects of rainfall fluctuations are more important than the final steadystate capacity vegetation towards which the system is ‘attracted’. The key outcomes of this work may be summarized thus: 1. For the three rainfall series and for the selected types of abandonment, the simulated sediment yield varies with rainfall amount and intensity as well as with the percentage and pattern of abandonment. 2. The sediment yield decreases with increasing vegetation cover following abandonment but, for very high rainfall intensities, the sediment yield for the vegetated catchment is not very different from the bare catchment. 3. In order to reduce the sediment yield, the best strategy is to select the most vulnerable areas for abandonment, while using those less vulnerable for agriculture. 4. The highest sediment yield is produced when areas close to the channel are not abandoned. These results have significant implications for the management of Mediterranean catchments. They show

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that control over the time-space pattern of deforestation/land abandonment can be used to reduce land degradation even at times of climate change and call for a more thoughtful approach to basin management from the forestry industry. They further invite caution about the assertion that deforestation is the main agent of Mediterranean land degradation by demonstrating the wide variety of possible responses.

Grazing Grazing is so frequently blamed for land degradation in the Mediterranean as to merit special attention in this chapter and some examples of blame are given below. There is need for some caution in adopting a singleprocess explanation because of the variable character of the grazing/land degradation interaction in space and time, the relatively poor understanding of the phenomenon, and the tendency of managers and politicians to seek a simple explanation for what appears to be a ‘quick-fix solution’ that can be easily implemented such as a limited and fixed ‘carrying capacity’. There are five main positions in the debate. First there is the view of those who consider that grazing by domestic animals is among the major causes of land degradation, with goats being singled out for their predilection for woody forage (e.g. Thirgood 1981, 1988). Tsouris (1985) claims that the pastoral economy had a much greater impact on deforestation in the Mediterranean than agricultural clearances. Second, there are animal ecologists (e.g. Papanastasis 2004) who claim that, since livestock numbers have been drastically reduced, in some parts of the Mediterranean undergrazing has resulted and this has led to the piling up of flammable biomass and hence increased fire risk (Chapter 19). Third, as with the forestry debate, there were quite different initial conditions in different parts of the Mediterranean, with livestock grazing spreading at different rates in the Holocene, being introduced 3,000 years later in the west than in the east. And they did not find a grazing-free environment, since wild herbivores were already there. In Crete, for example, wild fauna had arrived by the Middle Pleistocene (750,000–128,000 years BP) (Rackham and Moody 1996). Fourth, Papanastasis (1998) makes the point that Mediterranean ecosystems were affected not only by the grazing per se, but by the whole livestock economy, which involved other activities too such as shepherding, nomadism and transhumance, sheltering, milking, cheese-making, and especially burning. The latter (still practised today) was used to create grazing and suppress unwanted, chiefly woody vegetation in range-

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lands. The fifth approach to the debate attempts to establish the detailed process dynamics of the plant-animal interaction. One example is the work of Briske and NoyMeir (1998) that showed that several ecological scales must be considered when evaluating the response of vegetation to grazing to minimize incomplete and inaccurate conclusions. Ultimately, biomass reduction by consumption and by selective grazing leads to loss of cover and thereby to increased erosional susceptibility, but there are other effects too. The soil can be compacted, leading to crusting and/or higher levels of runoff and erosion. Regular animal routes (pathways) can become concentrations of overland flow that lead to gullying. Concentrations of livestock around watering points can produce the socalled ‘piosphere’ effect, whereby badly degraded lands form at these point locations with vegetation decreasing towards the water points. Some authors argue that the increased addition of manure has the reverse effect. The overall balance appears to depend on cattle density and raises the question of rapid change after a threshold has been crossed. Another important effect is the change in composition of the vegetation caused by the selective grazing of the more palatable plants. Recently it has been shown that erosion can develop and spread from erosion hotspots (Thornes 2005). There are two main processes at work here: the origin of the bare spots amongst otherwise fairly uniform vegetation cover and the spreading out of erosion from these hotspots. The original spots may be random or uniform and circular or linear and are often attributed to cattle concentrating for resources (food or water) or to random variations in topography or soil quality. Hotspots have been reported in the dry north-western sector of the island of Lesvos, Greece, but they have also been reported beyond the Mediterranean too, for example in the sub-humid lands of the Eastern Cape, South Africa (Kakembo 2003) and in the much wetter and cooler Icelandic heathland (Gisladottir 2004). A study of the emergence of such bare spots reveals that they can arise from uniform vegetation covers when there are positive feedbacks in the relationship between moisture, vegetation growth, and animal dispersion. They produce spotty patterns in the vegetation biomass cover that need no special initial patterning to develop and evolve. This approach (Cronhjort 2002) also shows that the phenomenon can lead to apparent movement of island-like patches of vegetation through time and that the emergence is related to climatic variations. A recent study from the European Union GeoRange Project has used the concept of ‘resistance’ surfaces. In this approach, the available forage for animals is

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mapped from satellite data and the barriers to movement (such as steep rocky slopes or poor grazing) are used to identify the optimal grazing location. This, taken in turn with soil and hillslope hydrological information, facilitates mapping and prediction of erosion-susceptible areas due to grazing. This methodology appears to offer great possibilities for future grazing management. Both approaches operate on representations of the landscape as cellular maps and hence are suitable for the application of cellular automaton modelling techniques that are rapidly being developed in Spatial Ecology (Dieckmann et al. 2002). As the impact of global warming becomes more acute in the Mediterranean, the threat to grazing difficulties seems likely to increase. Indeed, in a study of stocking capacity resilience to changing climate, Köchy et al. (2008) have shown through hierarchical modelling that grazing can result in greater reductions in 3,000

productivity as climate becomes drier. (Thornes and Brandt 1994) showed that the competition between vegetation and erosion leads to an oscillatory behaviour in both that arises from the intrinsic dynamics, prompted by external forcing (Figure 20.9). This is represented in a ‘boom and bust’ cycle. The forage increases dramatically in good years but then is unable to survive in dry years and drops to very low levels, thus leading to severe erosion. This can happen without cyclical rainfall series but it seems likely that global warming might progressively destabilize the grazing/soil erosion interaction and that cyclical behaviour is a symptom of this instability in the system. Ongoing research based on the BIOME model by Zhang at King’s College London (personal communication) has compared the land use of 1981–90 to 2040–50 with respect to climate change, as indicated by accepted European climate models. It shows a marked reduction (15%) of xerophytic

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wood/scrub and a dramatic reduction (nearly 96%) of cool grass/shrub. The benchmark (1981–90) was validated against Olson’s European remotely sensed land use data. The results should be taken as a warning of real and serious changes ahead. The destabilization of the grazing/soil erosion interactions would push the system into greater instability than that which already exists. In recent years, important lessons have been learned from the grazing management practices of sub-Saharan savanna rangelands, which should be relevant to Mediterranean pastoralism. Scoones (1989) has traced the history of colonial rangeland management practices and Behnke et al. (1993) have identified their weaknesses. The main shortcoming was the rigid application of the notion of carrying capacity and the development of elaborate algorithms produced to estimate it. The main problem is that a single number is being used to represent a complicated, complex, and highly variable interaction (the indicator fallacy). Second, it implies a stable situation in which management is all that is needed to maintain the balance. Third, the approach allows little insurance in times of global climatic or economic instability and pastoral livelihoods are undermined. The preference now is for a more flexible, opportunistic approach that provides the farmer with a wider range of alternative methods of adjustment to extreme factors. A more effective management strategy calls for a better recognition of the thresholds and instabilities in the grazing systems and a better capacity to predict the trajectories of change following perturbations, whether physical or human. Another problem is the tendency to reach for the ‘overgrazing explanation’ when the real explanation may only be revealed by careful analysis of the case history (Rowntree et al. 2004).

Land Degradation: Prospects for Improvements So far in this chapter it has been shown that land degradation is a real and present problem in the Mediterranean. It has been demonstrated that it is a complicated (multifaceted) problem and a complex one (there are multi-stable states to be considered). Moreover it is a ‘wicked’ problem in the sense of Ludwig (in Pritchard and Sanderson 2002)—in other words it involves a host of traditional academic disciplines or social perspectives. This leads to the adoption of fixed and often rigid positions. Given this situation, it has been difficult to evaluate the two prevailing shibboleths, grazing and forest clearance, so they have grown in strength and persist up to the present day.

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In the light of these ongoing debates, the expected crisis from global warming in the coming fifty years, and the sustained pressure on rural development, what can be done? What are the options for addressing land degradation in the Mediterranean? Evidently there are no quick and easy fixes, but future actions must take into consideration the changing conceptual and technical environments.

Technical Advances The major change in information brought about by remote sensing and computational capacity has greatly enhanced our ability to monitor the quality of the environment and track its changes through time. The CORINE Programme, for example, illustrated the huge potential of land use analysis. Perhaps greatest of all is the ability to assess the quantity and quality of vegetation and soils and especially the possibility of estimating the status of animal forage. Coupling this information with Eagleson’s theory of ecological optimality should provide an alternative strategy for estimating the annual water budget and its transformations across the Mediterranean. Also, with the increased availability of accurate and high-resolution digital elevation models, erosion modelling has moved into a new and more promising phase when coupled with vegetation cover data. Given the immense effort required to produce erosion estimates from physically based computer simulation models, they have been restricted to relatively small areas. New data and more recent models (such as the TOPMODEL, WEPP, and MEDALUS models) seem to offer fresh possibilities for obtaining estimates of erosion. They will, however, continue to need good-quality field data for calibration and validation.

Conceptual Advances Conceptually, administrators have to face up to the new paradigm of ecological complexity and its implications for understanding land degradation as outlined above. These conceptual changes have been extended, synthesized, and integrated into a metaphor (or model) for adaptive cycles of change in ecological and economic systems called ‘panarchy’ (Holling et al. 2002). The essential proposition is that there is an evolutionary cycle, comprising four phases (Figure 20.1). Starting from a collapse of the system caused by internal dynamics or external forcing, there is a phase of exploration and development as the system redevelops and reorganizes itself (the r-phase). As the system’s resources increase through time, the internal structure

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of the system becomes progressively stronger, coupled, and conservative, leading to a conservative phase (the K-phase) with an optimization of available resources. These are the traditional r- and K-phases of logistic growth, dominated by the r- and K-type species of earlier ecologists. Holling et al. (2002b) have added two further phases. The release (or Ÿ) phase in which the system collapses with a large loss of resources before wholesale reorganization of the system occurs, and the restructuring (·) phase that restarts the cycle. The parallels with the boom/bust cycle modelled by Thornes and Brandt (1994) and illustrated in Figure 20.9 are to be expected from the latter’s use of the Lotka–Volterra competitive-logistic modelling strategy (Lotka 1925; Volterra 1926). In this study it was argued that the unstable cycle results from the Lotka– Volterra approach. The r- and K-parameters were tuned by Mediterranean vegetation and climate values and the model has further been used to test the stability of the grassland-tree interaction in South Africa and the emergence of hot spots in homogeneous vegetation covers (Thornes 2005). Later in their treatise, Holling et al. (2002b) show how the basic metaphors can be extended to develop cascaded cycles at different time and space scales, with leakage of resources and constraining behaviour between different levels (Figure 20.10). This new and rich paradigm introduces the full range of complexity characteristics to evolving ecosystems and nicely removes the difficulties of the conventional

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(Clementsian) model of succession. It has important implications for ecological evolution in a land degradation context, even though it is neither fully worked out, nor is the panarchy theory fully tested. The proposed cycle is intrinsic and gives rise to periodicity through the internal dynamics and is constrained by changes in the external forcing agents that may themselves be periodic or simply stochastic. We can hypothesize that the degradation we see in the Mediterranean today is not the same as that observed by the Greek philosophers, but merely a different, later adaptive cycle. From the management point of view, the paradigm prompts the question of where we are in the cycle. The problem is to identify the critical thresholds of stability and instability, separating the zone of trajectories that lead either to complete vegetation cover or to absolutely bare soil and catastrophic erosion (Thornes 1985). Once identified, managers can decide if the system is on the brink of change when tiny adjustments might produce massive effects. When the panarchy paradigm is coupled to the new spatial ecology through the hypercycle (Dieckemann et al. 2000), real advances are in prospect. In the new spatial ecology, two- and three-dimensional spatial structures emerge and result in apparent movement of vegetation and erosional patches across the landscape. The other big shift, conceptually, is from the soil conservation approach, stressing technical fixes and practical measures, to the soil husbandry movement that stresses an integrated effort coupled to traditional

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methods and public involvement. The Australian Land Care experience (Roberts 1992) that arose from the National Soil Conservation Programme catalysed rural action and provided a framework for community land care action. In 1989 it was recognized that ‘[t]here is a need for a major national integrated program, which incorporates a local, co-operative, self-help approach and the development of individual property plans and increased assistance to state and local governments within an overall framework of policy guidance. The active co-operation of landholders is pivotal to the success of the program’ (Roberts 1992: 19). This is clearly different from the heavily bureaucratic approach enshrined in the Water Framework Directive for catchment management in Europe (Chapter 8). The new technical approach to conservation lays emphasis on engineering structures that reduce surface runoff rather than attempting to reduce soil erodibility. The main shift has been towards the increased use of vegetation in soil conservation. This is now known as bioengineering. An excellent example is the study by Quinton et al. (2002) of the use of bioengineering principles for mitigating desertification. This work examines the effect of above- and below-ground plant properties on infiltration, runoff generation, and soil erosion. It then uses this information and the available literature to suggest a number of species that might prove useful in a regeneration programme. They include an extensive list of plant species, their habitats, and ecological/engineering properties. Another example of a fresh approach to conventional wisdom of the mitigation of degraded areas is the study by Margaris and Koutsidou (2002). The study, from the Greek island of Chios, is on the interaction between fire and grazing. The result shows that, on the Aegean Islands, it is clear that the productive capacity of the pastures is not enough to support the current number of animals. The proposed solution is not simply a call to reduce a mythical carrying capacity, but rather the elimination of the practice of deliberately starting fires. On the basis of their deep understanding of the system, they seek in this way to alter the form of stock farming from intensive to extensive, by subsidizing stock farmers to replace the grazed biomass by imported animal feeds. The results of a cost study show that the economic benefits of these activities could be more effectively deployed to reduce the fire hazards. Land degradation is ultimately a problem of sustainability and stability. UNEP’s Agenda 2000 required nations to move towards sustainable land use. Although Mediterranean countries, north and south, are moving

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slowly in this direction, there is now a wide recognition of the general stability concepts that underline the search for solutions. We are entering a new, more progressive era in the land degradation problem, equipped with new paradigms and new mind-sets that bode well for the future.

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Fantechi, R. and Margaris, N. S. (1986), Desertification in Europe. Reidel, Dordrecht. Francis, C. F. and Thornes, J. B (1990), Runoff hydrographs from three Mediterranean vegetation cover types, in J. B. Thornes (ed.), Vegetation and Erosion. John Wiley & Sons, Chichester, 363–85. Freschi, A. (1998), Desertification in Italy: activities at national and regional level, in Burke and Thornes (1998: 77–110). Garcia Ruiz, J. M., Lasanta, T., Ruiz-Flano, P., Marti, C., Ortigosa, L. M., and Gonzalez, C. (1994), Soil erosion and desertification as a consequence of farmland abandonment in mountain areas. UNEP Desertification Control Bulletin 25: 27–33. Gisladottir, G. (2004), Ecological disturbance and soil erosion on grazing land in south-west Iceland. in A. Conacher (ed.), Land Degradation. Kluwer, Dordrecht, 109–26. Godron, M., Guillerm, J. L., Poissonet, P., Thiault, M., and Trabaud, L. (1981), Dynamics and management of vegetation, in F. de Castri, D. W. Goodall, and R. L. Specht (eds.), Mediterranean-Type Shrublands. Elsevier, Amsterdam, 317–42. Gomes, J. L., Arrue, J. L., Lopez, M. V., Sterk, G., Richard, D., Gracia, R., Sabre, M., Gaudichet, A., and Frangi, J. P. (2003), Wind erosion in a semi-arid agricultural area of Spain, Catena 52: 235–56. Goodess, C. and Palutikof, J. (2002), Local and regional responses to global climate change in south-east Spain, in N. A. Geeson, C. J. Brandt, and J. B. Thornes (eds.), Mediterranean Desertification: A Mosaic of Processes. John Wiley, Chichester, 247–69. Grove, A. T. and Rackham, O. (2001), The Nature of Mediterranean Europe: An Ecological History. Yale University Press, New Haven. Gunderson, L. H. and Holling, C. S. (eds.) (2002), Panarchy: Understanding Transformations in Human and Natural Systems. Island, Washington. Holling, C. S. and Gunderson, L. H. (2002), Resilience and adaptive cycles, in Gunderson and Holling (2002: 25–63). and Ludwig, D. (2002a), In quest of a theory of adaptive change, in Gunderson and Holling (2002: 3–24). and Peterson, G. D. (2002b), Sustainability and panarchies in Gunderson and Holling (2002: 63–103). Holmgren, N. and Scheffer, M. (2001), El Niño as a window of opportunity for the restoration of degraded arid ecosystems. Ecosystems 4: 151–9. IPCC (Inter-Governmental Panel on Climate Change) (2001), Third Assessment Report. Kakembo, V. (2003), Factors affecting the invasion of Pteronia (blue bush) onto hillslopes in Ngqushwa (formerly Peddie) District, Eastern Cape. Ph.D. thesis, Rhodes University, South Africa. Kirkby, M. J. (1980), The problem, in M. J. Kirkby and R. P. C. Morgan (eds.), Soil Erosion. John Wiley & Sons, Chichester, 1–16. Köchy, M., Mathaj, M., Jeltsch, F., and Malkinson, D. (2008), Resilience of stocking capacity to changing climate in arid to Mediterranean landscapes. Regional Environmental Change 8: 73–87. Kosmas, C., Danalatos, N. G., Lopez-Bermudez, F., and RomeroDiaz, M. A. (2002), The effect of land use on soil erosion and land degradation under Mediterranean conditions, in N. A. Geeson, C. J. Brandt, and J. B. Thornes (eds.), Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 57–81.

Langbein, W. B. and Schumm, S. A. (1958), Yield of sediment in relation to mean annual precipitation. American Geophysical Union Transactions 39: 1076–84. Lavee, H., Poesen, J., and Yair, A. (1997), Evidence of high efficiency water harvesting by ancient farmers in the Negev desert. Journal of Arid Environments 35: 341–8. Le Houérou, N. (1981), The impact of man and his animals on Mediterranean vegetation, in F. De Castri, D. W. Goodall, and R. F. Specht (eds.), Mediterranean Type Shrublands. Elsevier, Amsterdam, 479–521. Lesschen, J. P., Kok, K., Verburg, P. H., and Cammeraat, L. H. (2007), Identification of vulnerable areas for gully erosion under different scenarios of land abandonment in Southeast Spain. Catena 71: 110–21. Llamas, M. R. (1988), Conflicts between wetlands conservation and groundwater exploitation: two case histories in Spain. Environmental Geology and Water Sciences 11: 241–51. Lotka, A. J. (1925), Elements of Physical Biology. Williams & Wilkins, Baltimore. Margaris, N. S. and Koutsidou, E. (2002), Landscape protection from grazing and fire, in N. A. Geeson, C. J. Brandt, and J. B. Thornes (eds.), Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 83–92. Märker, M., Angeli, L., Bottai, L., Costantini, R., Ferrari, R., Innocenti, L., and Siciliano, G. (2008), Assessment of land degradation susceptibility by scenario analysis: A case study in Southern Tuscany, Italy. Geomorphology 93: 120–9. Mtimet, A. (1999), Atlas des Sols Tunisiens. Ministry of Agriculture, Republic of Tunisia. Mulligan, M. (1998), Modelling the complexity of land surface response to climatic variability in Mediterranean environments, in M. G. Anderson and S. M. Brooks (eds.), Advances in Hillslope Processes. John Wiley & Sons, Chichester, ii. 1099–150. Nafaa, R. and Watfeh, A. (2002), Holocene and actual degradation of the environment in the Mamora Forest (Morocco). International Journal of Anthropology 15: 263–70. Naveh, Z. and Dan, J. (1973), The human degradation of Mediterranean landscapes in Israel, in F. de Castri and H. A. Mooney (eds.), Mediterranean Type Ecosystems, Springer, New York, 373–90. Obando, J. A. (2002), Modelling the impact of land abandonment on regeneration of semi-natural vegetation: A case study from the Guadalentin, in N. A. Geeson, C. J. Brandt, and J. B. Thornes (eds.), Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 419–29. Papanastasis, V. P. (1998), Livestock grazing in Mediterranean ecosystem: an historical and policy perspective, in V. P. Papanastasis and D. Peter (eds.), Ecological Basis of Livestock Grazing in Mediterranean Ecosystems, EUR18308EN. Office for Official Publications of the European Communities, Luxembourg, 5–10. (2004), Traditional vs. contemporary management of Mediterranean vegetation: the case of the island of Crete. Journal of Biological Research 1: 39–46. Price, P. N. (1998), Actions taken by Moroccan national agencies to combat desertification, in Burke and Thornes (1998: 113– 49). Pritchard Jr., L. and Sanderson, S. E. (2002), The dynamics of political discourse in seeking sustainability, in Gunderson and Holling (2002: 147–32).

Land Degradation Pueyo, Y. and Alados, C. L. (2007), Effects of fragmentation, abiotic factors and land use on vegetation recovery in a semiarid Mediterranean area. Basic and Applied Ecology 8: 158– 70. Quinton, J., Morgan, R. P. C., Archer. N. A., Hall, G. M., and Green, A. (2002), Bioengineering principles and desertification mitigation, in N. A. Geeson, J. B. Thornes, and C. J. Brandt (eds.), Mediterranean Desertification: A Mosaic of Processes and Responses. John Wiley & Sons, Chichester, 93–105. Rackham, O. and Moody, J. A. (1996), The Making of the Cretan Landscape. Manchester, Manchester University Press. Reynolds, J. F. and Stafford Smith, D. M. (2002), Global desertification: do humans cause deserts? in J. F. Reynolds and D. M. Stafford Smith (eds.), Global Desertification: Do Humans Cause Deserts? Dahlem University Press, Berlin, 1–21. Riksen, M., Brouwer, F., and de Graaff, J. (2003), Soil conservation policy measures to control wind erosion in north-western Europe. Catena 52: 309–27. Roberts, B. (1992) Land Care Manual. University of New South Wales Press, Sydney. Rodda, J. C. (1976), Basin studies, in J. C. Rodda (ed.), Facets of Hydrology. John Wiley & Sons, Chichester, 257–98. Rowntree, K., Duma, M., Kakembo, V., and Thornes, J. B. (2004), Debunking the myth of overgrazing and soil erosion. Land Degradation and Development 15: 203–14. Salvucci, G. D. and Eagleson, P. (1992), A Test of Ecological Optimality for Semi-arid Vegetation. Report 335, Ralph M. Parsons Laboratory, MIT, Cambridge Mass. Scoones, I. (1989), Economic Development and Zimbabwe’s Communal Areas. Pastoral Development Network Paper 27b, Overseas Development Institute, London. Thirgood, J. V. (1981), Man and the Mediterranean Forest. Academic Press, New York. (1988), Goat grazing in Cyprus. Unasylva 40: 59. Thomas, D. S. G. and Middleton, N. J. (1994), Desertification: Exploding the Myth. John Wiley & Sons, Chichester. Thornes, J. B. and Brandt, C. J. (1994) Erosion-vegetation competition in a stochastic environment undergoing climatic change, in A. C. Millington and K. Pye (eds.), Environmental Change

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This chapter should be cited as follows Thornes, J. B. (2009), Land degradation, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 563–581.

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21

Water Resources Jean Margat

Introduction The geography of natural water resources in the Mediterranean basin cannot simply be reduced to the study of water inputs, water distribution, and the pattern of runoff-generating precipitation determined by climate and relief—although these are, of course, fundamental controls (Margat 1992; Benblidia et al. 1996). Any consideration of basin-wide water resources also needs to consider a range of territorially determined factors affecting water resources. These include: (1) the nature of surface and underground flows, which depends on river basin and hydrogeological characteristics; (2) the natural storage capacity of lakes and aquifers and their role in regulating flows, and any losses from these stores which reduce the resulting flows; (3) the existence of favourable conditions for water management and exploitation such as suitable sites for dam construction and the productivity of aquifers, as these factors dictate accessibility to water resources and the production costs; (4) the natural quality of the water, its vulnerability to pollution and its capacity for self-purification; (5) any constraints imposed for reasons of environmental conservation, which may effectively exclude a proportion of water reserves from the category of exploitable resources.

The Geography of Mediterranean Water Resources It is important to appreciate that each of these factors influences the assessment of water resources in a given area and each factor has its own geography (Margat 1997; Margat and Vallée 1999a). In spite of the broad similarities in climate and landscape between the different parts of the Mediterranean basin, there are

considerable variations between regions that impact upon the availability of water resources. Many of the factors affecting water resources cited above are subject to a similar degree of variation (Grenon and Batisse 1989; Chapter 8) and these are discussed in turn below.

The Uneven Distribution of Inputs Marking the transition between the temperate climate of Europe and the aridity of North Africa and the Near East, the Mediterranean climate contains wide variation, and this is reflected in a highly uneven distribution of rainfall (Benblidia et al. 1996; Margat and Vallée 1999a ; Chapter 3). For example, moving from one extreme to another, average annual rainfall ranges from more than 3,000 mm in parts of the Dinaric Alps to less than 50 mm in Libya. There is, first and foremost, a contrast between north and south, but east–west variations are also noticeable. In the west, the Mediterranean regions of Spain and France receive rather less rainfall than some coastal areas of the Maghreb. In the east, in Anatolia and the Levant, precipitation inputs decrease steeply from the coast towards the interior and in Egypt and Libya the Sahara desert extends right up to the coastal zone. These patterns are highlighted in Figure 21.1 which shows the basin-wide distribution of potential annual runoff or effective precipitation.

Blue and Green Water Precipitation inputs in the Mediterranean basin give rise to two types of water resource: (1) the surface and groundwater flows, or ‘blue water’, which make up the conventional natural renewable resources, and (2) the soil moisture, or ‘green water’, which can be utilized by agriculture (Falkenmark and Rockstrom 2004). Both types of water resource are unevenly distributed across the Mediterranean basin. The average yearly runoff that

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Jean Margat

> 500 mm/yr 200 – 500 100 – 200 50 – 100 < 50

0

500 km

Fig. 21.1. Distribution of potential annual runoff (effective precipitation) in the Mediterranean basin.

forms the domestic resources of each country ranges from more than 1 million m3 per km2 in some mountainous areas in the north, to less than 10,000 m3 per km2 in the arid zones of the south, or, in terms of local rainfall inputs, it ranges from more than 1,000 mm per year to less than 10 mm (Figure 21.1). The annual average amount of ‘green water’ depends on the real level of evapotranspiration from cultivated land and varies from 700 mm at the most in the north to less than 100 mm in the south, or from 7,000 m3 to less than 1,000 m3 per hectare per year. This means that irrigation is necessary almost everywhere.

Drainage Basin Size The storage zones that collect and retain blue water are very fragmented and this resuslts in water resources being distributed between a large number of drainage basins. Such fragmentation demands a large network of independent measurement and management approaches. The Mediterranean basin is dominated by short and steep river channel systems than drain small catchments (Chapters 8 and 11). The topography of the region dictates that the headwaters of most rivers lie above 500 m in the upland zone and the watershed of river systems that drain to the Mediterranean Sea is close to the coast in many countries (Woodward 1995)

as shown in Figure 21.2. In fact, only twenty-one river basins that drain to the Mediterranean Sea cover an area of more than 10,000 km2 , and only three (the Nile, Rhône, and Po basins) cover more than 50,000 km2 (Figure 21.2). Some basins are also divided by national borders, meaning that their resources are shared by several countries. This is especially true of the Iberian and Balkan peninsulas and the Near East, the most notable case being that of the Nile basin, which is shared by ten countries, but for the most part lies outside the Mediterranean zone (Howell and Allan 1994; Woodward et al. 2007). Similarly, a great number of aquifer systems, generally limited in extent and closely linked to surface watercourses, are found mainly in alluvial or karstic and sometimes volcanic contexts. These are often formed in highly compartmentalized hydrogeological structures, with the sole exception of the sedimentary basins of the northern Sahara which impinge only slightly on the Mediterranean basin. The quantity of renewable water resources contained in these basins is, however, minimal.

Fluvial Runoff and Flow Regimes The small size of many Mediterranean drainage basins and the low level of rainfall inputs in less humid areas combine to reduce average stream flow volumes

Water Resources

585

Mean annual discharge > 1,000 m3 s-1 Mean annual discharge 100 – 1,000 m3 s-1 Other permanent rivers Seasonal or ephemeral streams

0

500 km

Fig. 21.2. The Mediterranean drainage network and river basins.

(Chapter 8). Excluding the Nile, only twenty-five watercourses (and all of them are found in the northern Mediterranean) have average flows greater than 100 m3 s−1 (Figure 21.2). Large rivers are the exception in the Mediterranean basin, where only the Rhône, the Po, and the Nile have average flow rates of over 1000 m3 s−1 (3.5 km3 per year). These rivers owe their high volumes to the fact that their catchments extend beyond the climatic zone of the Mediterranean into the Alps or equatorial Africa (Chapter 8).

Flow Regimes Owing to the generally low levels of reservoir storage, there is a great deal of variation in flow rates throughout the Mediterranean basin—both between seasons and between years. The low flows of summer are always far lower than the annual mean flow—commonly being less than one-tenth, or even one-hundredth of monthly winter flows and, where there is no contribution from groundwater, they can be zero (Chapter 8). Overall, for the whole of the Mediterranean basin, the low flows of regular surface resources amount to only a fifth of average total flows, and do not exceed 100 km3 per year. However, this irregularity of flow is not present

to the same extent in all regions and it becomes more accentuated the further south one goes, as increased aridity aggravates these hydrological contrasts (Margat 1998a ).

Surface Flows and Transmission Losses Mediterranean watercourses do not retain all the water they collect as they suffer considerable transmission losses due to evaporation or consumption by riparian vegetation (Servat et al. 2003). Their outlet flow, especially at the point of discharge into the sea, can be significantly lower than the sum of inputs into their basin. As a proportion of inputs, these losses—which are amplified by human consumption (especially through irrigation)—can, for example, reach 35 per cent in the Ebro basin, 18 per cent in the Medjerda basin in Tunisia, and 64 per cent over the entirety of the Nile basin prior to the construction of the Aswan High Dam (the real flow rate of the Nile into the Mediterranean has since been further reduced). It is therefore possible to underestimate potential resources if they are only calculated solely on the basis of river discharge measurements taken in the lower reaches of river basins.

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Natural Storage Capacity

Water Quality

The natural water reserves of the Mediterranean basin (both on the surface and in the groundwater store) are, in total terms, quite large. The Alpine lakes of the Po and Rhône basins alone have a capacity of 220 km3 , and one could add the reserves contained in glaciers to this figure (Chapter 12). However, these reserves have only a moderate regulatory effect on river flows, as the levels, and therefore the water reserves, of lakes and alluvial aquifers vary little, and karstic aquifers can empty quickly (Chapter 10). Some of the very deep, multi-level limestone karsts around the Mediterranean basin have very large storage capacities. These systems can cope with large interannual recharge variations and they often support exploitation under high pumping rates (Bakalowicz et al. 2008). The Saharan aquifers on the southern and eastern border of the Mediterranean basin (Egypt, Libya, Tunisia) that were mentioned earlier store large quantities of water, but this is essentially ‘fossil’ water with a very low rate of renewal and therefore does not contribute significantly to the regulation of river flows. No permanent watercourses exist in these arid regions and the groundwaters constitute a nonrenewable resource (Margat 1998b).

The natural quality of surface water and groundwater varies across the Mediterranean and often fluctuates over time both in respect of chemical composition and biotic content. In the north, quality is usually satisfactory for most uses, most notably for purification into drinking water, and for most aquatic ecosystems. Nitrate pollution of groundwater has been reported in many areas and is largely due to the use of nitrogen-rich fertilizers. On the island of Majorca nitrate concentrations in groundwater of up to 700 mg l−1 have been reported (Candela et al. 2007). In the catchments in the south, water quality often falls in association with falls in water quantity, so low quality frequently goes hand-in-hand with scarcity of resources. The greater water treatment costs—especially for drinking water production—that are thus necessitated to a lesser or greater extent, further accentuate the inequality in terms of available resources between the north and the south. The most common natural water quality issues, both in the north and the south, are the hardness of groundwater in carbonate-rich areas (which are found throughout the basin) and the often high turbidity of surface waters. However, it is the often high salinity of perennial water that reduces freshwater resources in southern countries. In Tunisia, for example, 26 per cent of surface water, 90 per cent of phreatic groundwater, and 80 per cent of deep-pumped groundwater has a salinity of more than 1.5 g per litre. However, it is difficult to classify the quantities of moving water in each basin in terms of quality, since the quality of water flowing within a single watercourse tends to decline as it goes downstream. The degree to which Mediterranean water, especially groundwater, is naturally protected against pollution or other threats to its quality (in other words, its vulnerability), also varies across the region. Degradation caused by humans is more widespread and affects a larger quantity of Mediterranean water than degradation by natural processes.

Water Resources and the Impact of Human Activities The Influence of Land Use The flow regimes of Mediterranean river basins, especially the smaller ones, are very sensitive to the state of the soil and of plant cover in the upstream catchment and hence to the impact of human activities that can alter these characteristics. In general, this tends to lead to an increase in the irregularity of their flows (Chapter 8). Their much diminished base flow rates greatly reduce their capacity for self-purification and increase their vulnerability to pollution. Prat and Munné (2000) examined the controls on water quantity and quality in the Congost stream of north-east Spain. The high demand for water from agriculture, the urban environment, and industry means that channel flows are effectively maintained by effluent returns from sewage plants and little or no dilution takes place from natural stream flows. Such lack of dilution has severely degraded the ecological productivity of the channel zone. Since most of the groundwater stores of alluvial plains, particularly in coastal areas, are easily accessible and exploitable, they are also vulnerable to the risk of overexploitation. They are also often poorly protected against pollution threats as explained below.

The Uneven Distribution of Water Resources Water Resources by Country The natural renewable water resources of the Mediterranean basin—in other words the sum of surface water and groundwater flows—total about 600 km3 in an average year and data for each country are presented in Table 21.1 (Margat and Vallée 1999a ). However, according to their size and their climate, countries within the basin can possess internal resources that

TABLE 21.1. Water resources in the Mediterranean basin by country and continent Countries and territories in the Mediterranean basin

Natural renewable resources (km3 /year) in an average year Internal resources 1 Surface watera

2 Groundwaterb

3 Overlapc

4 Totald 1 + 2−3

Spain France Italy Malta Slovenia Croatia Bosnia-Herzegovina Serbia-Montenegro FYR Macedonia Albania Greece

26.6 63.5 172.5 0 4.1 10.0 14.0 15.5 5.4 25.7 55.5

10.4 32.0 43.0 0.05 3.0 ∼9 ∼2.0 ∼2.0 ∼1 6.2 10.3

∼9 31.5 33 0 2.9 ∼1 ∼2 ∼1.5 ∼1 ∼5 7.8

28 64 182.5 0.05 4.2 ∼18 14 ∼16 5.42 26.9 58

Total North

392.8

118.95

—

417.1

Turkey Cyprus Syria Lebanon Israel Palestine: West Bank Gaza

External constraints: proportion of natural resources

6 Total 4+5

of external origin: index of dependence (%)

13.65 0 0 1 14.8 16.25

28.35 72.5 191.3 0.05 4.2 32 14 16 6.42 41.7 74.25

1.2 10.6 3.9 0 0 42.7 0 0 14.9 36.8 16.8

23.8e

440.9e

5 External resourcese 0.35 8.5 8.8 0

Â

20.0 0.4 2.4 3.2 0.45

∼15 0.3 ∼2 2.5 0

66 0.8 5 4.8 0.63

3.45 0 0.96 0 0.38

69.45 0.8 5.96 4.8 1.01

6.6 0 12.3 0 38.0

0.07 0

0.5 0.05

0 0

0.57 0.05

0 0.01

0.57 0.06

0.0 16.7

77.8

2.8e

80.6e

37.4

−

Egypti Natural Res. Real Res. Libya Tunisia Algeria Morocco

0.3 0.3 0.2 2.9 11.64 4.9

0.5 8j 0.6 1.15 1.33 1

0 7.5j 0.1 ∼0.35 ∼1 ∼0.9

0.8 0.8 0.7 3.7 11.97 5

84 55.5 0 0.32 0.03 0

85 56.3 0.7 4.02 12 5

Total South

19.94

4.58

—

22.17

55.5e

77.7e

Overall

483.4

160.9

—

517.1

e

82

599.2e

0.1 0.9 0 0 92.2 0 97 82.5 100 0 2.7

21.4 40 110 0.015g 2.2 9 7 8 3 13 29

Present production capacity of non-conventional water (km3 /year) Regeneration of used water available for reuse 0.2 —

Reusable drainage water

∼0.2

0.0016

— — — — —

Desalination

0.13 — 0.058 0.022 —

97 0 97–8

— — —

— 0.013

96

242.6

61.0 0.7 4.6 4.1 0.18

Total East

flowing into neighbouring countries (restrictions on freedom of use) (%)

Estimated exploitable resourcesf (km3 /an)

0.3 0 29h 10.6 1.0 100 0

40 0.54 4 2.18 0.75 0.53 0.05

—

Â

— 0.011

98

0.034

0.02 0.27

94 96

0.007 0.085 0.03

0.027

00 99

Â

48.05 98.8 98.6 0 8.0 0.21 0

0 0 0 0 2.2 0.75

49.5 0.65 1.8 6 3.4

0.7 0.07 0.026

12.6 98 97

0.27 0.054 0.06 0

99 00 90

61.35 352

a Total runoff; b Groundwater recharge; c Overlap: groundwater flows collected or fed by watercourses. Equivalent to the difference between the value of column 2 and the groundwater that flows directly into the sea or across borders, or evaporates in depressions in arid zones; d Total without double counting; e Surface or groundwater flow from a neighbouring country, whether or not it adjoins the Mediterranean. The external resources of each country are not simply added to obtain the subregion figure because of sharing between countries. These figures are therefore free of double counting, and refer only to water brought in from non-neighbouring countries. The figures for the total resources in subregions and in the entire basin (column 6) take this into account; f According to criteria appropriate to each country. These figures are therefore not homogeneous; g Malta: water that is exploitable without upsetting the freshwater/saltwater balance; h Syria: on the basis of the natural outflow of the Orontes into Turkey; i Egypt: the distinction between natural and real resources is crucial here on account of the importance of inflows from the Nile; j Egypt: additional groundwater recharge by irrigation in the Nile valley and delta: 7.5 km3 /year of ‘secondary resources’. Source: National sources compiled by le Plan Bleu (2002).

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Jean Margat

Fig. 21.3. Natural renewable and exploitable water resources per country in the Mediterranean basin (annual means).

amount to literally billions of cubic metres per year for the better off, or just millions of cubic metres for the most deprived. Thus, the three countries with the greatest water wealth in the Mediterranean basin, namely France, Italy, and Turkey, together possess well over half (60%) of the internal water resources of the whole basin. By contrast, the least well-off parts of the basin (Malta, Gaza, Cyprus, and Libya) have only 3 per cent of this total (Figure 21.3 and Table 21.1). To these internal resources should be added the external resources brought by cross-border rivers, as the Mediterranean basin extends to some countries that do not adjoin the Mediterranean Sea itself. While these resources have a secondary role in Western Europe and the Maghreb, they are more important in south-eastern Europe (Albania, Croatia, Greece) and are, naturally, crucial in Egypt, where the proportion of water resources that originate from abroad is overwhelming (98.8%). The predominance of external resources in the south-east of the Mediterranean is essentially due to the contribution of the Nile to Egypt as shown in Table 21.2. The differences in the exploitability of resources further deepen the divisions between countries.

Water Resources by Population Calculating the water resources available per head of present-day populations provides deeper insights into the levels of water wealth or poverty in the Mediterranean basins of each country. The distribution is again highly uneven (Figure 21.4), ranging from extreme water poverty (less than 100 m3 per year in Gaza, less than 200 m3 per year in Israel, Libya, and Malta), to abundance (more than 10,000 m3 per year in Albania, Bosnia-Herzegovina, Croatia, Slovenia, and SerbiaMontenegro). At the time of writing, 114 million Mediterranean people (45% of the total population of the basin), spread TABLE 21.2. Key figures on internal and external natural and exploitable water resources for the three main regions of the Mediterranean basin Subregion of the basin North (Europe) East (Near East) South (Africa) TOTAL

Natural internal % Real external Total resources resources resources (km3 /year) (km3 /year) (km3 /year) 417 78 22 517

81 15 4 100

24 3 55.5 82.5

441 81 77.5 599.5

%

73.5 13.5 13 100

Water Resources

589

Fig. 21.4. Natural renewable water resources and real exploitable water resources per inhabitant (in 2000) of each country in the Mediterranean basin.

over eight countries, live under the water poverty level of 1,000 m3 per year in natural water resources (annual average) per inhabitant. Below this level, tensions generally arise between requirements and resources, particularly in countries where irrigation is necessary. In six countries or territories (with a total of 48 million inhabitants) resources per head fall below even the level of absolute ‘penury’ of 500 m3 per year: Israel, Libya, Malta, the Palestinian territories, Tunisia, and Algeria. Furthermore, several of these countries are either already using almost the entirety of the exploitable resources in their sector of the Mediterranean basin, or have exceeded this level, thereby necessitating the use of imports from outside the basin (Israel, Libya). This inequality becomes even more acute when differences in levels of socio-economic development are considered, as it is often in the poorest countries that resources are scarcest and therefore most expensive to exploit (see Oki and Kanae 2006).

Evaluating and Exploiting Water Resources Irregular inflows and outflows, limited monitoring networks for both surface water and groundwater,

reductions in the streamflow of many watercourses, exchanges between interdependent reserves of surface water and groundwater, together with the growing impact of human activities on water regimes, make it very difficult to obtain the necessary hydrological data that would form the basis of an evaluation of the resources of natural and renewable water in the Mediterranean basin. Gathering these data requires that a larger number of observations be made over longer periods than is the case in other regions, and also that appropriate corrections be applied to stream discharge measurements. As perennial surface water flows that may be exploited by direct withdrawal generally only constitute a small proportion of surface water resources, it is often necessary to implement regulatory schemes in order to manage these resources. The distribution of dams across the different areas of the Mediterranean basin is uneven, and the sites of these dams vary with regards to their profitability, ease of maintenance, and their location in relation to points of water use. The potential for controlling water supply throughout the year by using very large capacity reservoirs is very limited. Few of the reservoirs so far constructed in the Mediterranean have a capacity approaching 1 km3 , and there is only

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one giant reservoir in the basin: the Aswan High Dam reservoir (Lake Nasser) on the Nile (164 km3 ), but this is an unusual example and is way beyond the southern limit of the Mediterranean region. In the whole of the Mediterranean basin, the total quantity of water held in artificial reservoirs (not counting Lake Nasser) currently stands at just 66 km3 . However, the potential capacity of reservoirs that could still be created tends not to be included in these calculations, as the feasibility criteria of these projects are complex and subject to change. There are many potential sites for small reservoirs in various northern and southern (especially in Morocco, Tunisia, and Algeria) countries, but their usefulness as regulatory sources and their useful lifespan are both reduced by the risk of sedimentation (Lahlou 1988). Sources of groundwater vary in size, and their accessibility and exploitability depend upon the depth and productivity of the aquifers in question. Furthermore, they often, particularly in the north, feed perennial surface water flows. The complications that arise from this interdependence place serious limits on their exploitability, and may also generate conflicts between users. In several countries, a proportion of groundwater flows directly into the sea via coastal or submarine karst springs that are difficult to tap; over the Mediterranean as a whole, these flows account for 17 per cent of the basin’s groundwater (Chapter 10; Bakalowicz et al. 2003; Fleury et al. 2008). The often considerable difficulties involved in water management limit the proportion of truly usable natural resources. The irregularity of an often overwhelming proportion of the water flows in the region necessitates the construction of water management systems that vary greatly in cost and efficiency. In addition to this, the decline in efficiency that accompanies the expansion of these systems and of water exploitation schemes causes costs to rise. The efficiency of artificial reservoirs is reduced by the high evaporation rate to which their contents are subject (from 0.8 m per year to more than 2 m per year). In Algeria, for example, evaporation from reservoirs stands at between 1.3 and 2.2 m per year. In Egypt, evaporation causes the loss, on average, of 10 billion m3 per year from Lake Nasser and this is equivalent to about 12 per cent of the controlled flow of the Nile. The position of water management installations in relation to points of use—such as towns and irrigation schemes—often leads to high transport and pumping costs. Using criteria that are as much economic as social in nature, it can be argued that not all Mediterranean water can be counted as a resource because people are either unable or unwilling to exploit part of that resource. A proportion of non-perennial surface water,

most notably in the case of heavy flooding, cannot be controlled, either because of the lack of a suitable site for a dam, or because to control it would necessitate the construction of installations that would be impractical on grounds of cost or environmental impact. A proportion of perennial surface water and groundwater is also unusable because it is needed for on-site applications or for the preservation of wetland ecosystems (Chapter 9).

Competition for Water Resources Water, like any resource, is the object of rivalries and even of conflicts between the growing needs of users, all the more so where it is in short supply. Competition between abstractors of groundwater and surface water users frequently arises, often being driven by ignorance of the interdependence between these two components of a single resource. Groundwater abstraction can cause springs to dry up and reduce the flow rates of streams in dry periods, while dams can prevent the recharge of aquifers. The largest area of competition is between users in towns and their rural counterparts. The provision of drinking water in cities, generally a priority in social and political terms, is set against irrigation, the area of use normally hardest hit by demand management measures due to its predominant position among the various sectors of current water consumption. All the users and river basin managers also need to consider the ecological requirements of the aquatic system itself (Prat and Munné 2000). The Mediterranean environment was previously richer in perennial streams and wetlands than it is today (Chapter 9). Over time human activities have caused these to recede, whether unintentionally, as a consequence of water withdrawal, or voluntarily, through the drainage of marshes judged to be unhealthy or the reclamation of cultivable land. In Spain, in recent years, for example, 40,000 hectares of wetlands have been lost—three-quarters of which were linked to intensively exploited groundwater resources. Contemporary ecological concerns, in particular with regard to the preservation of the aquatic environment and wetland zones, which are becoming increasingly rare, have necessitated the designation of ‘restricted outflows’, which can impose local restrictions, varying according to national environmental policies, on water management and withdrawal. In this way a proportion of surface and groundwater is excluded from the total quantity of exploitable resources. The flashy discharge peaks of Mediterranean river systems mean that flooding is a major problem, and

Water Resources

this hazard is magnified by the increasing concentration of population and human activities in high-risk zones (Chapter 18). The risks are often underestimated due to complacency or the short-term interests of those occupying the land in question. However, in the north and in the south, a string of recent disasters, which are not as ‘natural’ as some might argue, serve as a reminder that water in the Mediterranean is not only a resource but also a potential danger. In many Mediterranean basins the prevention of water-related risks is one of the principal objectives of water management.

Threats to Mediterranean Water Resources Environmental changes caused by the people of the Mediterranean are reducing the renewal rate and the quality of water resources (Bethemont 2000; Jeftic et al. 1996). Water renewal cycles are threatened by various types of land use that have been transforming the Mediterranean environment for centuries. This process of transformation has accelerated in the twentieth century. Rampant urbanization (particularly in coastal areas) and deforestation in the south, along with the cessation of slope maintenance through terraced agriculture in the north, can all lead to enhanced runoff volumes and sediment transfer (Chapter 8). In combination with the engineering of stream channels, these changes to river basins have made stream flows less predictable and, it can be argued, have reduced the total quantity of reliable water resources. Groundwater levels are very sensitive to the effects of over-exploitation, especially when abstraction is carried out by disparate groups who do not coordinate their activities, and whose objectives are short-term in nature. This is a particular problem in coastal areas, where the equilibrium between groundwater and seawater can easily be disrupted, leading to virtually irreversible salt water contamination of coastal aquifers. This has happened in most Mediterranean countries, but the effects can be strongly conditioned by local geological and hydrological factors, as Calvache and PulidoBosch (1997) have shown for three coastal aquifers in southern Spain. The quality of surface waters and groundwaters can be degraded by various sources of pollution. A range of pollution types can be identified, including point sources of effluent returns, industrial waste, and accidents, and diffuse sources from the use of fertilizers and pesticides in intensive agriculture, or poor waste management. Groundwater is more vulnerable in this respect as,

591

while cleaning up groundwater pollution is more timeconsuming, any loss of groundwater quality may impact upon perennial surface watercourses. Groundwater-fed springs are commonly the best source of potable drinking water in much of the Mediterranean and the bottled spring water trade is an important industry. Surface water stored in reservoirs, as well as in natural lakes, is also under threat from eutrophication triggered by land use and climatic factors (e.g. Reed et al. 2008; Chapter 9). Pollution can wipe out a proportion of water resources, rendering them unusable, or push up drinking water purification costs to a prohibitive level. Pollution is not the sole preserve of industrialized northern countries, in which it is undoubtedly more prevalent, but also dealt with more effectively. It is becoming more common in the south, where underdeveloped purification systems and preventive measures increase the likelihood of pollution events, and the consequences of these events are aggravated by the scarcity and frequently low natural quality of perennial water resources.

The Sustainability of Water Management Measures The high suspended sediment loads found in floodwaters in the Mediterranean basin, especially in the south (Chapter 8), causes the silting-up of reservoirs to be a particular water resource problem. It shortens the useful lifespan of the reservoirs as regards flow regulation, in spite of the enormous ‘sleeping reserves’ that they were designed to hold. In the south, the loss rate of useful reservoir capacity currently runs at between 0.5 and 1 per cent per year, sometimes higher: 0.5–2 per cent in Algeria, where the lifespan of average-capacity reservoirs is between thirty and fifty years; 0.5 per cent in Morocco, where the reduction of regulatory capacity attributable to silting is currently equivalent to a loss of irrigation potential of between 6,000 and 8,000 hectares per year; and, finally, 1–2.5 per cent in Tunisia. Algeria’s reservoirs have already lost 25 per cent of their total initial capacity, while Morocco’s had lost 8 per cent (800 million m3 ) of theirs by 1990, with certain reservoirs already being half-filled by silt (Lahlou 1988; Chapter 8). Since only a limited number of sites are suitable for the construction of dam and reservoir installations, and some of these have already been developed, the development and silting up of all of these reservoir sites is to be expected in the longer term, probably before the end of the twenty-first century. Efforts at prevention through the stabilization of hillslopes and the installation of

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sediment traps can at best stall this process, but cannot prolong reservoir life indefinitely in such high sediment yield environments. A significant decline in the amount of water resources available from these forms of regulatory management is thus inevitable. As regards the exploitation of the non-renewable resources that are provided by a number of large aquifer reservoirs in several southern countries (Libya, Tunisia, Algeria), especially in the Saharan regions on the outer rim of the Mediterranean basin, the necessarily limited duration of exploitation is governed, like any drilling operation, by the rate of extraction that has been selected, and will at best be in the order of fifty years. In addition to this, the quality of the water extracted from these aquifers can be degraded through contamination by saline water even before the reserves are exhausted, further shortening their exploitable lifespan.

The Threat of Climate Change The water resources of the Mediterranean basin are not protected from the effects of any climate change caused by the intensification of the greenhouse effect over the twenty-first century (Margat 2004). In the south, the risk of the climate becoming more arid is inescapable, and this would have the dual effect of reducing water resources and increasing water requirements as a result of higher evaporation rates and lower rainfall. In the north, the danger is rather that the climate will become more extreme, wetter in the winter and drier in the summer, and more unpredictable (Chapter 3). This would also have consequences for patterns of water use, possibly raising water requirements in the summer, while increasing the risk of flooding. Water resources are not simply a raw material provided by nature to be exploited until it can no longer renew itself. They are vulnerable to the impact of the various effects of their exploitation and of water use. They are also vulnerable to the effects of the various lifestyles and activities of the human population. Mediterranean water resources are naturally fragile, but this fragility is increased by intensive exploitation and by the activities of Mediterranean people, which have radically altered the natural water budget. The greater the use made of these resources, the more vulnerable this use becomes to unpredictable natural factors such as droughts and falls in water quality, and to direct human-induced pressures. The water resources of today are not, it would seem, quite what they were in the past even if the standards of evaluation have themselves evolved. The evaluation

of water resources in the Mediterranean basin needs in particular to combine the hydrological data necessary to predict their future potential with feasibility studies of their sustainable exploitation. All this should be coupled with an understanding of their sensitivity to environmental factors that takes into account the limits imposed by the need for their conservation. Two opposing tendencies could lead to the revision of calculations of exploitable water resources in the Mediterranean basin. On the one hand resources may increase as, in the face of diminishing availability, the pressure of demand forces up the ceiling of the water management and exploitation costs that are judged to be acceptable (within the limits, nonetheless, of countries’ differing economic capacities and the physical limits of natural resources). Alternatively, resources may drop, owing to an increasing awareness of the limits imposed on exploitation by the need for environmental conservation. The final outcome may vary according to the countries in question, differing water and environmental policies, and changes in the overall situation.

A Diminishing Resource At present, pressures on resources vary, but are often intense. In the Mediterranean basin, the geography of water use and hence of water requirements is no less varied than that of water resources, although there are certain dominant trends such as the predominance of irrigation and the increasing requirements of cities and of the tourist industry. Throughout history, demand for water has been met exclusively by the management and exploitation of natural water resources. As demand has increased, so these management measures and withdrawals have intensified. The levels of channel diversions, control of irregular flows by dams and reservoirs, and exploitation of groundwater by tapping or pumping vary, but in several countries these are already high. The growing scarcity of sites with development potential, the reduction in the flow of perennial streams, and also the intensive exploitation of groundwater, with all its undesirable consequences, are key problems for water resource planners in the region. The scarcer resources become the more sensitive they are both to quantity-related pressures resulting from withdrawals and quality-related pressures caused by the recirculation of used water. Although the different aspects of these pressures throughout the Mediterranean basin result in a somewhat varied picture, the pressure on resources has already reached high levels in most areas. In the Mediterranean basin as a whole,

Water Resources TABLE 21.3. Annual water withdrawal volumes for the three main regions of the Mediterranean basin Subregion in the basin

North (Europe) East (Asia Minor, Near-East) South (Africa)

Total draw-off (km3 /year)

Proportion of natural renewable resources (average exploitation index) (%)

90 18 69

19 22 88

177 billion m3 of water are withdrawn and used annually and this amounts to a quarter of the average level of natural resources, and certainly a larger proportion of exploitable resources. The total volume of annual water withdrawals is shown in Table 21.3 for the European, African, and Asian sectors of the Mediterranean basin. However, withdrawal levels and pressures on resources differ considerably according to subregions and even more so according to countries and river basins (Figure 21.5). In several countries, the quantity of withdrawals from the basin is approaching or indeed has equalled the quantity of resources in an average year, and therefore

593

threatens to exceed it in a dry year. The exploitation indexes for renewable water calculated per country in the basin still show considerable variation, but nonetheless exceed 50 per cent in eight countries (Table 21.4). They therefore indicate that tense situations exist, at least at the local level and according to local circumstances. In some countries the maximum level has been reached, approaching or sometimes even exceeding 100 per cent. This may indicate a loss of equilibrium or the deliberate but unsustainable recourse to non-renewable resources (Libya), or the fact that a proportion of the resources is being used more than once, such as in the collection and reuse of used water (Israel), or the recirculation of drainage water (Egypt). Final natural resource consumption figures, which represent the quantity of water that is withdrawn and then not replaced after being used, are also high in Mediterranean countries, where two factors help to raise them: (1) the relative importance of high-consumption agricultural uses and (2) the high proportion of used water from towns, industries, and tourist developments that, as a consequence of coastal population concentration, is discharged into the sea and therefore reduces the amount of water returned to streams or groundwater.

75 – 100% 50 – 75% 25 – 50% 0– 25%

0

500 km

Fig. 21.5. The exploitation index of natural renewable water resources across the Mediterranean basin. This is the ratio of withdrawals to the sum of internal and external resources (see Table 21.4).

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Jean Margat TABLE 21.4. Present-day pressures on water resources in the Mediterranean basin

Countries and territories in the basin

Spain France Italy Malta Slovenia Croatia Bosnia-Herzegovina Serbia-Montenegro FYR Macedonia Albania Greece

Date Total

Surface water

Groundwater

1997 1999 1998 1997–8 1996 1996 1995 1995 1996 1995 1997

17.94 16.67 42 0.025 0.03 0.19 0.1 0.8 1.85 1.40 8.7

14.71 14.72 31.6

10.25 5.0 18 0.004f

0.02 0.04 0.02 0.5 1.65 0.77 5.14

3.23 1.95 10.4 0.025 0.01 0.15 0.08 0.3 0.2 0.63 3.56

89.7

69.17

20.54

41.2

1997 2000 1997 1996 1999 1994 1996

11.1 0.295 3.85 1.3 1.12 0.13 0.13 17.93

10.25

7.68

10.1

1995–6 1998 1996 2000 1998

59.5 2.0 2.27 2.9 1.9

55g 0.1 1.15 2.0 1.7

4.5 1.9 1.12 0.9 0.2

35.1 1.3 1.1 1.7 0.9

Total North Turkey Cyprus Syria Lebanon Israel Gaza West Bank Total East Egypt Libya Tunisia Algeria Morocco Total South OVERALL

Final consumption (km3 /year)a

Withdrawals (km3 /year)

68.57 176.2

Â

6.3 0.15 2.85 0.9 0.05

 Â

59.95 139.4

4.8 0.145 1.0 0.40 1.07 0.13 0.13

Â

0.15 0.1 0.3 0.8 0.6 6.0 Mean %

6.0 0.2 1.6 1.0 1.1 0.1 0.07

Indicators of pressure on water resources (%) Exploitation indexb

Final consumption indexc

84 42 38.2 162d 1.5 2 1.4 10 60 10 29

48 13 12.5 67f 0.15 0.8 0.6 3.4 25 4.5 17

18.7 28 55 96e 80 120 260 25

Mean %

21.7 96 233 69 41 56

8.6 15 41 40 47 110 108 9.4 12.2 110 200 42 23 34

8.62

40.1

Mean %

87.9

51.4

36.83

91.4

Mean %

27.5

14.3

Notes: a Net consumption by use plus used water not returned to continental waters (discharged into the sea); b Exploitation index: annual withdrawals/average yearly flow of total natural renewable resources in %; c Consumption index: annual final consumption/average yearly flow of total natural renewable resources in %; d Relates to resources that can be exploited without breaking the freshwater/saltwater balance; e Syria: relates to real resources; f Malta: includes returns of used water of non-conventional origin (desalination); g Egypt: gross withdrawals, including recovery of drainage water. Sources: National and international figures compiled by le Plan Bleu (2002).

At present, final resource consumption over the whole of the Mediterranean basin is probably near to 100 billion m3 per year, or 70 per cent of the total draw-off, 55 per cent of which occurs in the southern (40.1 km3 per year) and eastern (10.1 km3 per year) countries (Table 21.4). Naturally, these quantity-related pressures are accompanied by the impact of discharges of used water upon water quality as well as that of other sources of pollution. In an average year, around 20 billion m3 , a significant proportion of which is not purified, are released into the continental waters of the Mediterranean basin. While they are less industrialized than northern countries, southern countries have scarcer water resources, and so may suffer the effects of

this pollution to a greater degree. In Egypt, for example, where industries discharge 550 m3 of waste water per year (57% of all by-products) into the Nile, the number of deaths attributable to illnesses caused by water pollution in the Nile have, according to the WHO, risen from 19,395 in 1979 to 48,458 in 1987. The efforts undertaken to improve sanitation, the purification of used water, waste management infrastructure, and the general prevention of pollution have not, by and large, been as extensive as those put into expanding water supplies. This is also the case in other parts of the world, but the consequences in the Mediterranean basin are more serious because its water resources are scarcer, and have greater demands placed upon them.

Water Resources

Temporal Trends in Water Resource Pressures The rate of increase of these pressures on water resources in the Mediterranean has varied, and continues to vary. The available national statistics do not allow a reliable historical account of water withdrawal to be established, not even for the period following the middle of the twentieth century. Notwithstanding the variation in their quality, the figures for the last twenty-five years revealed by these statistics are significant (Margat and Vallée 1999a). In Mediterranean countries as a whole, withdrawals have risen overall by 50 per cent, although this varies according to region, with increases of nearly 30 per cent in the north, where growth has slowed, of more than 60 per cent in the south, and a doubling of withdrawals in the east. The differences between individual countries are also considerable and will involve: (1) a tripling of withdrawals in Libya; (2) a near-doubling in Algeria, Spain, France, Syria, and Turkey; (3) increases of more than 50 per cent in Greece, Lebanon, Morocco, and Tunisia; (4) moderate growth in Egypt, and stabilization in Italy, where withdrawals are very high; (5) low growth or stabilization in Israel and Cyprus, where the ceiling of resources has been reached. In the Mediterranean basin, specific withdrawal histories are difficult to establish, but the general tendencies would seem to be largely convergent and exploitation figures have typically risen in the same proportions in many areas (Margat and Vallée 1999a, b). The pressures on resources will continue to grow in the twenty-first century, although to a lesser extent than most national and international projections have suggested (e.g. Vision Mediterraneenne de 2000). While future demand might be moderated by measures to improve the economy of water use, it will still, with a few exceptions, essentially be covered by the exploitation of natural resources. According to the studies carried out as part of the Plan Bleu project (2003), water demand and the withdrawals that it necessitates could grow in total across the Mediterranean basin, following current trends, by about 20 km3 per year between 1995 and 2025. Such an increase would give an overall total of 211.9 km3 per year by 2025 (Table 21.5). This amounts to a rise of just 11 per cent with considerable differences between subregions: a 4 per cent drop in the north, and rises of 20 per cent in the south and 45 per cent in the east. The highest growth is predicted to occur in the Mediterranean parts of Libya (+65%), Turkey (+60%), and Algeria (+50%). These trends would further widen the disparity between north and south.

595

TABLE 21.5. Water demand predictions for 2025 in the three regions of the Mediterranean basin Subregions in the Mediterranean basin North East South TOTAL

Trend projections of total water demand in 2025 (km3 /year) 86.2 27.8 97.9 211.9

Generally, then, it is in the countries where pressure on resources is already high that it is likely to increase the most. According to these projections of current trends, the exploitation indexes for natural renewable resources will be above the 50 per cent mark in the Mediterranean basins of nine countries in 2025, and will be approaching or will have already exceeded 100 per cent in Egypt, Israel, and Gaza. When applied to exploitable resources, these indexes would rise to 50 per cent in thirteen countries, while approaching or exceeding 100 per cent in eight countries. Water penury threatens the twenty-first century. Even according to the relatively optimistic hypothesis that natural water resources are sustainable (with the effects of climate change not being felt before the middle of the twentyfirst century); structural water penury would still be expected within a generation across a substantial part of the Mediterranean basin as a result of increased pressure on resources. Two further projections that show rising levels of water penury support this conclusion: first, there will be a drop in the level of water resources per inhabitant as a result of projected demographic expansion. Per-inhabitant water resources (average renewable resources) of below 1,000 m3 and 500 m3 are, respectively, accepted indicators of water poverty and absolute water penury. Second, given the decline in per capita resources predicted to occur between now and 2025, we are left with the prospect of 164 million Mediterranean people in eight countries living in a situation of water poverty (compared to 114 million at present). Six of these countries, with a combined population of 67 million people, will be in a situation of water penury (compared to 48 million at present). By 2025 more than half the population of the Mediterranean basin will therefore be in a situation of stress or penury in relation to natural water resources, let alone in relation to the resources that can actually be utilized. There will be a significant reduction in the net availability of water (resources remaining after withdrawals and recirculated water have been taken into account) in terms both of quantity and quality. This availability will

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Jean Margat

decline at a more or less rapid rate as a function of final consumption, and will tend to disappear completely in those countries and places where it is already very low (Malta, Israel, Gaza, and Libya). In this way, we can begin to sketch out the geography of present and future water penury in the Mediterranean basin. This reveals the same kinds of disparities as those found in studies of water resources and the pressures upon them. It can be said with certainty that the countries or territories that are under the greatest threat are: Malta, Gaza, Israel, and Libya, where structural penury is already present and then Cyprus, Tunisia, Syria, Algeria, and Egypt, where occasional instances of penury will become more frequent, and regional structural penury will become the norm.

Conclusions In common with other parts of the world (see Pimentel et al. 1997), naturally occurring water resources are distributed unevenly in the Mediterranean basin, while the ease with which they can be exploited to meet demand also varies markedly. These are major contributing factors to the disparity of development possibilities and cost levels in the provision of water supplies in the Mediterranean basin. These costs will, in general, increase, although once again unevenly and rising most steeply where they are already high. This will further handicap the development of the southern and eastern countries—especially in irrigated agriculture— more than that of northern countries. This will distort competition in the Euro-Mediterranean free-trade zone scheduled to begin in 2010. In an analysis of water scarcity and food import in six Mediterranean countries (Algeria, Egypt, Israel, Libya, Morocco, and Tunisia), Yang and Zehnder (2002) argue that water scarcity is a rigid limiting factor to food production and that food imports are imperative for compensating water resource deficiency. This situation increases the potential for conflicts of use, whether between sectors of use (principally between urban communities and agricultural irrigation), between the different regions of a country (upstream–downstream rivalry, regional opposition to water-transfer projects), or between countries forced to share resources such as in the Iberian Peninsula, the Balkans, the Near East, and the Nile basin (Maury 1990). Levels of water resources and water availability vary between neighbouring territories or countries, and this opens up the possibility of creating transfer schemes in order to alleviate situations of water

penury, whether between regions within a single country (these are already planned or in operation in Spain, France, Greece, Israel, Egypt, Libya, Tunisia, Cyprus, and Morocco), or between different countries, creating an international water trade. Indeed, projects between France and Spain (Catalonia), Albania and Italy, Turkey and the Near East have been proposed. The geography of water resources represents a key control on development and biodiversity in the Mediterranean basin and the pressures to which they are now subjected are expected to increase during the course of the twenty-first century. This poses major challenges for river basin planners across the region and for the implementation of the Water Framework Directive in the European Mediterranean. Given the population and climate change predictions for the region, the geography of present and future situations of water penury will become increasingly important in the economic geography and geopolitics of the Mediterranean world.

References Bakalowicz, M., Fleury, P., Jouvencel, B., Promé, J. J., Becker, P., Carlin, T., Dörfliger, N., Seidel, J. L., and Sergent, P. (2003), Coastal karst aquifers in Mediterranean regions: a methodology for exploring, exploiting and monitoring sub-marine springs. Tecnologia de la Intrusion de Agua de Mar en Acuiferos Costeros: Paises Mediterraneos. IGME, Madrid. El Hakim, M., and El-Hajj, A. (2008), Karst groundwater resources in the countries of eastern Mediterranean: the example of Lebanon. Environmental Geology 54: 597–604. Benblidia, M., Margat, J., and Vallée, D. (1996), L’Eau en région méditerranéenne—Water in the Mediterranean Region. Conférence Euroméditerranéenne sur la gestion de l’eau, Marseilles, November 1996; new edn. 1997, Blue Plan, Sophia Antipolis. Bethemont, J. (2000), La question de l’eau en Méditerranée. Revue de l’économie Méridionale, 48/191: 179–90. Calvache, M. L. and Pulido-Bosch, A. (1997), Effects of geology and human activity on the dynamics of salt-water intrusion in three coastal aquifers in southern Spain. Environmental Geology 30: 215–23. Candela, L., Wallis, K. J., and Mateos, R. M. (2007), Nonpoint pollution of groundwater from agricultural activities in Mediterranean Spain: the Balearic Islands case study. Environmental Geology 54: 587–95. Falkenmark, M. and Rockstrom, J. (2004), Balancing Water for Humans and Nature: A New Approach in Ecohydrology. Earthscan, London. Fleury, P., Bakalowicz, M., de Marsily, G., and Cortes, J. M. (2008), Functioning of a coastal karstic system with a submarine outlet, in southern Spain. Hydrogeology Journal 16: 75–85. Grenon, M. and Batisse, M. (eds.) (1989), Futures for the Mediterranean Basin: The Blue Plan. Oxford, Oxford University Press.

Water Resources Howell, P. P. and Allan, J. A. (eds.) (1994), The Nile: Sharing a Scarce Resource. Cambridge University Press, Cambridge. Jeftic, L., Keckes, S., and Pernetta, J. (1996), Climate Change and the Mediterranean. Edward Arnold, London, ii. Lahlou, A. (1988), The silting of Moroccan dams, in M. P. Bordas and D. E. Walling (eds.), Sediment Budgets. IAHS Publication 174: 71–7. Margat, J. (1992), L’Eau dans le bassin méditerranéen. Situation et prospective, Blue Plan 6. Éditions Economica, Paris. (1997), Cadre Géographique des Ressources en Eau dans la Région Méditerranéenne. Rapport au congrès international: L’acqua nei paesi mediterranei – Problemi di gestione di una risorsa scarsa. Consiglio Naz. delle Ricerche/Naples, 4–5 December 1997. CNR/IREM. Il Mulino, Bologna, 31–52. (1998a ), Sécheresses et ressources en eau en Méditerranée. Rapport à la conférence sur la politique de l’eau en Méditerranée, Valencia Espagne 16–18 August, 1998: Session Gestion des sécheresses. CR édités par le Réseau Méditerranéen de l’Eau, Madrid. (1998b), Les Eaux souterraines dans le bassin méditerranéen. Resources et utilisations. Documents du BRGM 282. Blue Plan– BRGM, Orléans. (2004) L’Eau des Méditerranéens. Situation et perspectives. Rapport Techniques du PAM, à paraître. and Vallée, D. (1999a ), The Mediterranean in Figures. Water Resources and Uses in the Mediterranean Countries. Figures and Facts. Blue Plan, Sophia Antipolis. (1999b), Vision méditerranéenne sur l’eau, la population et l’environnement au XXe siècle Mediterranean Vision for Water, Population and the Environment in the 21st Century. MEDTAC, document pour le IIe Forum mondial de l’eau de la Haye. Global Water Partnership, Conseil Mondial de l’Eau. Blue Plan, Sophia Antipolis.

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Maury, R.G. (1990), L’Eau dans les pays méditerranéens de l’Europe communautaire. Études méditerranéennes 15. Centre Interuniversitaire d’Études Méditerranéennes, Poitiers. Oki, T. and Kanae, S. (2006), Global hydrological cycles and world water resources. Science 313: 1068–72. Pimentel, D., Houser, J., Preiss, E., White, O., Fang, H., Mesnick, L., Barsky, T., Tariche, S., Schreck, J., and Alpert, S. (1997), Water resources: agriculture, the environment, and society. BioScience 47: 97–106. Prat, N. and Munné, A. (2000), Water use and quality and stream flow in a Mediterranean stream. Water Research 34: 3876–81. Reed, J. M., Leng, M. L., Ryan, S., Black, S., Altinsaçli, S., and Griffiths, H. I. (2008), Recent habitat degradation in karstic Lake Uluabat, western Turkey: A coupled limnological– palaeolimnological approach. Biological Conservation 141: 2765–83. Servat, E., Najem, W., Leduc, C., and Shakeel, A. (2003), Hydrology of Mediterranean and Semi-arid Regions. IAHS Publication 278. International Association of Hydrological Sciences, Wallingford. Woodward, J. C. (1995), Patterns of erosion and suspended sediment yield in Mediterranean river basins, in I. D. L. Foster, A. M. Gurnell, and B. W. Webb (eds.), Sediment and Water Quality in River Catchments. John Wiley & Sons, Chichester, 365–89. Macklin, M. G., Krom, M. D., and Williams, M.A. J. (2007), The Nile: Evolution, Quaternary river environments and material fluxes, in A. Gupta (ed.), Large Rivers: Geomorphology and Management. John Wiley & Sons, Chichester, 261–92. Yang, H. and Zehnder, A. J. B. (2002), Water scarcity and food import: a case study for southern Mediterranean countries. World Development 30: 1413–30.

This chapter should be cited as follows Margat, J. (2009) Water resources, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 583–597.

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22

Air Pollution and Climate Jos Lelieveld

Introduction It has long been known that atmospheric pollutants can be hazardous to human health and ecosystems. This includes effects from episodic peak levels as well as the long-term exposure to relatively moderate concentration enhancements. Environmental issues related to air pollution include acidification, mostly by the strong acids from sulphur and nitrogen oxides, eutrophication by the deposition of reactive nitrogen compounds, the reduction of air quality by photo-oxidants and particulate matter, and the radiative forcing of climate by increasing greenhouse gases and by aerosol particles. Many air pollutants are photochemically formed within the atmosphere from emissions by traffic, energy generation, industry, the burning of wastes, and forest fires. The Mediterranean basin in summer is largely cloudfree, and the relatively intense solar radiation promotes the photochemical formation of ozone (O3 ) and peroxyacetyl nitrate (PAN); O3 being health hazardous at levels in excess of about 100 Ïg/m3 . Ozone is formed in the lower atmosphere as a by-product in the oxidation of reactive carbon compounds such as carbon monoxide (CO) and non-methane volatile organic compounds (NMVOC), catalysed by nitrogen oxides (NOx ≡ NO + NO2 ). In summer, notably the period from June to August, transport pathways of air pollution near the earth’s surface are typically dominated by northerly winds, carrying photo-oxidants and aerosol particles from Europe into the Mediterranean basin. Aerosol particles with a diameter of less than ∼10 Ïm (PM10 ) can have adverse health effects at a concentration of about 30 Ïg/m3 or higher. The fine mode particles (). The study domain, with a focus on the eastern Mediterranean region, as observed from the NASA SeaWiFS satellite in August 2001, is shown in Figure 22.2. The campaign was performed using Crete as an operating base to study the long-range transport and photochemistry of air pollution, including the effects on air quality and climate. Some of the main results have been presented by Lelieveld et al. (2002a). The foremost conclusion from the MINOS studies is that summertime air pollution not only adversely affects human health throughout the Mediterranean basin, but also that aerosol particles strongly influence the radiation energy budget in the area, perhaps conspiring with some of the regional climate consequences of global atmospheric change.

Climate and Meteorology The Mediterranean climate is characterized by humid winters with cyclonic storms, and warm, dry summers, with occasional extended drought periods (Bolle 2003; Chapter 3). The average north–south temperature

NOx 50.0 10.0 5.0 1.0 0.5 0.1

NMVOC 50.0 10.0 5.0 1.0 0.5 0.1

Fig. 22.1. European air pollution emissions, 2000, in thousand tonnes per 50 km grid cell (after Lövblad et al. 2004). NOx emissions in units of NO2 .

gradient across the basin, from the Alps to North Africa, is remarkably large, about 25◦ C. In some locations, e.g. along the Adriatic coast, precipitation is among the highest in Europe (∼1,000 mm per year) and in the mountains of the Balkan Peninsula it is typically >2,000 mm per year, whereas in much of North

Air Pollution and Climate

601

Fig. 22.2. Widespread aerosol haze in the Mediterranean basin, as observed from the SeaWiFS satellite during the MINOS campaign in August 2001 (© Orbimage, NASA).

Africa this can be more than an order of magnitude less. The Mediterranean weather is strongly influenced by the geographical positions of the Azores high and the Icelandic low, which can modify the mean westerly flow (Traub et al. 2003). When the high and the low are strongly developed, and therefore the meridional pressure gradient is strong, relatively moist air masses are transported to Europe. In the alternate case, the zonal flow is weak, and blocking weather systems prevail over central Europe. The interannually varying pressure gradient between the Icelandic low and the Azores high pressure systems is known as the North Atlantic Oscillation (NAO), an indicator of the inten-

sity of synoptic weather systems over the North Atlantic Ocean. Even though the NAO is a natural mode of climate variability, it can be influenced by anthropogenic climate change (Corti et al. 1999). It is linked to the general atmospheric and oceanic circulation systems, the latter being affected by the formation of cold bottom water in the Arctic Ocean and the influx of salty water from the Mediterranean Sea through the Strait of Gibraltar (Bolle 2003) (Chapter 2). By definition, a relatively strong pressure gradient between the Azores high and the Icelandic low yields a positive NAO index, a relatively weak gradient a negative NAO index. Although its influence has been studied primarily for the winter season,

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the NAO is associated with considerable monthly and interannual variability, and effects have been identified for all seasons. During the MINOS campaign in August 2001, the mean NAO index was 0.3. In summer, the InterTropical Convergence Zone (ITCZ) and the Azores high shift to higher latitudes. The Azores anticyclone, in combination with eastward moving low pressure systems over central Europe, lead to westerly flow in the lower and middle troposphere towards the Mediterranean basin. Spreading of the high pressure across central Europe in summer weakens the westerly flow, and European air is mostly transported to the basin by northerly winds, at right angles with the strong east–west pressure gradient near the surface. In the upper troposphere, air masses are transported in the westerlies from the North Atlantic Ocean (Figure 22.3). During spring and summer the Asian continent heats up, developing a heat low over northern India, Pakistan, and Iran. This causes the ITCZ to move north, generating the Indian monsoon with intense deep convection. In the upper troposphere, a semi-permanent anticyclone is located over the surface heat low, called the Tibetan high. The tropical easterly jet stream is an inherent feature of the Indian summer monsoon and the Tibetan high. It is a belt of strong easterly winds at the southern periphery of the upper tropospheric anticyclone. The tropical easterly jet, between 100 and 200 hPa pressure altitude, transports air from Asia over northern Africa and, aided by the upper tropospheric anticyclone over the Arabian Peninsula, towards the eastern Mediterranean (Traub et al. 2003). Under the influence of the subtropical jet stream and the westerlies, the air is subsequently transported back to the east over the Asian continent. During summer the Mediterranean basin is directly under the descending branch of the Hadley circulation, driven by deep convection in the ITCZ. In the upper troposphere, poleward moving air from the ITCZ is deflected by the Coriolis force. The resulting westerly flow reaches a maximum in the subtropical jet stream at about 40◦ N in summer, north of the Mediterranean Sea. As a result of subsidence, the region is characterized by dry and cloud-free conditions with high solar radiation intensity. Over land, convection can develop, which is generally not very deep at coastal locations. The air near the surface travels to the equatorial region as the synoptic northerly flow combines with the trade winds, carrying moisture evaporated from the Mediterranean Sea southward. In summer, teleconnections between the Mediterranean region, the Indian monsoon, and the Sahel rainfall regimes can be important, varying on an interannual

Upper troposphere

Middle troposphere

Boundary layer

500 km Fig. 22.3. Schematic of air flows during the MINOS campaign in August 2001. The arrows indicate atmospheric transport during approximately 2–4 days (after Lelieveld et al. 2002a).

time scale (Ward 1992). The atmospheric pressure at sea level in the eastern Mediterranean is anticorrelated with that in the Indian monsoon, mainly in the July– September period, while that in the western Mediterranean is positively correlated, with a maximum during September–November. The meridional wind component over the central and eastern Mediterranean basin is anticorrelated with that in the Indian monsoon. This means that a more active monsoon is connected with lower atmospheric pressure at sea level over the eastern

Air Pollution and Climate

part of the basin and higher pressure over the western basin, and stronger northerly winds over most of the Mediterranean. The northerly flow in the basin near the surface during summer, in combination with the lack of precipitation, promotes the long-range transport of pollutants from Europe (Luria et al, 1996). Air pollutants released in southern Europe may even reach the ITCZ over Africa on a time scale of 4–6 days (Kallos et al. 1998). Since the Mediterranean basin is largely surrounded by mountains, the local meteorology can be strongly influenced by orographical flows with strong diurnal cycles (Milláan et al. 1997). These flows can combine with land–sea breeze systems, so that relatively deep land-inward surface winds can be established covering distances up to 60–100 km. Since many pollutant sources are located in coastal regions, these circulations can carry air pollution into convective-orographical ‘chimneys’ that connect the surface flow to the return flow above, i.e. linking into the regional northerlies (Millán et al. 2005). Similarly, daytime orographical flows in the Alps can vent air pollution from the valleys into the lower free troposphere, where these air masses can be transported towards the Mediterranean (Henne et al. 2004). Further downwind over the sea, the signatures of these local circulation systems are evident as distinct pollution layers in the lower ∼4 km of the atmosphere, maintained over long distances owing to the stable stratification in the region. The following section addresses the long-distance air pollution transport that can be distinguished in the Mediterranean upper troposphere. This is followed by a discussion of the atmospheric chemistry and transport pathways near the surface, including the consequences for air quality and climate. The final section discusses the links between air quality, climate, and the water cycle in view of atmospheric changes expected in the twenty-first century.

Upper Tropospheric Pollution Plume In summer, the Tibetan high in the upper troposphere combines with the high over the Arabian Peninsula into a large anticyclone, which carries air from the Indian monsoon region over Africa to the eastern Mediterranean. The thunderstorm clouds in the ITCZ region over southern Asia can carry both moisture and air pollution from the surface to high altitudes on a timescale

603

of several hours, thus injecting the pollutants into the upper tropospheric anticyclone. During MINOS the impacts of this feature have been detected over Crete and the Aegean Sea, for example, by enhanced concentrations of CO, CH4 , and non-methane hydrocarbons in an extensive plume extending in altitude from the middle troposphere (6–8 km) up to the tropopause (∼15 km). Tracer-transport modelling studies indicate that the upper tropospheric pollution plume from South Asia was present for about 25 per cent of the time, whereas alternatively it was deflected to the south during which the Mediterranean upper troposphere is influenced more strongly by air mass transports from North America and the North Atlantic region (Lawrence et al. 2003). Air mass back trajectory calculations combined with global tracer transport modelling furthermore shows that the transport time of the South Asian plume, for example, from the polluted megacities in India to the eastern Mediterranean upper troposphere, is about 10–15 days (Figure 22.4; Traub et al. 2003). Within the South Asian plume, CH4 concentrations are relatively high, probably as a result of wetland and agricultural emissions, mostly from rice fields. Carbon monoxide and non-methane hydrocarbons are also enhanced (Scheeren et al. 2003). The relatively high concentrations of PAN and the products from biomass combustion (such as C2 H2 , CH3 CN, and CH3 Cl), point to the importance of biofuel use, agricultural waste burning, and other incomplete burning processes in the South Asian source region. By combining trajectory modelling and chemical analyses, east–west gradients of air pollutants in the upper troposphere over the Mediterranean basin become apparent, associated with air pollution sources in South Asia and North America, respectively. For example, the biomass burning products were significantly higher towards the east, whereas fossil fuel combustion products (including CO2 ) were significantly higher towards the west. Also HCFC-134a, a halocarbon used in air-conditioning systems in automobiles in North America, was clearly enhanced towards the western part of the basin. The chemical ‘age’ of air pollutants in the South Asian upper tropospheric plume has been estimated from the enhancement or emission ratio of biomass burning species (X) relative to CO, X/CO, which changes as a function of the travel time and the photochemical lifetime of the species involved. The calculated emission ratios have been compared with those measured downwind of India in the Indian Ocean Experiment (INDOEX) in 1999 (Scheeren et al. 2002), providing additional evidence of the South Asian origin of the air pollution over

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Days 60°N

16 14 12

40°N 10 8 20°N

6 4 2

0° 0 20°W

20°E

60°E

100°E

Fig. 22.4. Transport-time spectrum showing the period between release of air pollutants at the surface in northern India and arrival in the upper troposphere (200 hPa) in August 2001 (M. Traub, personal communication).

the Mediterranean in the upper troposphere as well as the plume travel time of 10–15 days. The importance of the Asian summer monsoon in pollution transport—note that this is the region with the strongest growing emissions in the world—is not limited to the upper troposphere, but can even extend into the lower stratosphere. Over the Mediterranean basin in summer the tropopause strongly slopes downwards, from about 16 km over North Africa to about 11 km over southern Europe. It appears that in summer the westerly wind in the polar front jet and the southerly wind at the eastern flank of the upper tropospheric anticyclone converge near the tropopause (Traub and Lelieveld 2003). The consequent acceleration of the flow increases the vertical and horizontal wind shear, creating a jet streak within the subtropical jet stream and deep tropopause folds. The strong wind shear leads to clearair-turbulence and mixing of the pollutants across the tropopause.

The residence time of the pollutants in the lower stratosphere is about 2 weeks (5–30 days) after which these substances are transported back into the troposphere. Transport is in an easterly direction from the Mediterranean over the Black Sea and Asia towards the Pacific Ocean. Hence the combination of the Asian monsoon convection and the unique meteorology in the upper troposphere causes South Asian emissions to reach the lower stratosphere in summer, and contribute to the efficient long-distance transport of air pollution.

Ozone Transport and In Situ Formation Although the upper tropospheric pollution plume from South Asia is not very rich in ozone, the 4–6-km thick layer in the middle troposphere beneath this

Air Pollution and Climate

plume contains surprisingly high O3 mixing ratios up to about 240 Ïg/m3 (typically 130–200 Ïg/m3 ). By using a chemistry-general circulation model, Roelofs et al. (2003) have analysed the photochemistry and transport phenomena involved. It was found that the upper tropospheric anticyclone influences the potential temperature and vorticity distributions over the Mediterranean basin such that air is efficiently transported downward and south-eastward. Thus a large fraction (∼80%) of the air mass trajectories in the westerly wind regime, ending over the Mediterranean between 4 and 8 km altitude, originate in the upper troposphere over the North Atlantic Ocean (Traub et al. 2003). Interestingly, these trajectories are strongly influenced by stratosphere–troposphere exchange associated with cyclonic activity over the North Atlantic Ocean. As a consequence, nearly 30 per cent of the tropospheric ozone over the Mediterranean basin in summer originates from the stratosphere. In winter this fraction may even be larger (Kentarchos et al. 2001). In spite of the importance of stratosphere–troposphere exchange for the middle troposphere, near the surface the ozone is almost entirely produced by in situ photochemical production. Owing to the stable stratification of the Mediterranean troposphere, exchange between the marine boundary layer and the free troposphere is suppressed. The model calculations suggest that about 90 per cent of the O3 near the surface is locally produced, of which about 80 per cent is caused by the photochemical conversion of anthropogenic emissions transported in the northerly flow from Europe. The remainder is from natural ozone sources. Note that a substantial part of the NMVOC acting as an O3 precursor is of natural origin. Some Mediterranean tree species appear to be particularly efficient emitters of highly reactive monoterpenes (Kesselmeier and Staudt 1999; Simon et al. 2005). The combination of these reactive NMVOC and NOx from traffic and industrial sources can give rise to a rapid build-up of ozone. Not surprisingly considering the northerly flow, the western Mediterranean is mostly influenced by emissions in Western Europe, while the eastern Mediterranean is most strongly influenced by emission in Eastern Europe. The depth of the boundary layer over the Mediterranean Sea in summer is about 0.8–1 km, over which O3 mixing ratios exceed 100 Ïg/m3 during most of the summer season, and the diurnal variation is only small, about 10 per cent. These relatively high mixing ratios in background air seem to be typical for the entire

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basin, as shown by measurements (Millán et al. 2000; Gangioti et al. 2001; Kouvarakis et al. 2000, 2002a; Nolle et al. 2002) as well as model results, which also indicate that Mediterranean ozone concentrations are among the highest in the Northern Hemisphere (Figure 22.5). Above the boundary layer the stable reservoir layer channels air pollution southwards. This pollution also originates in Europe, lofted to 1–4 km altitude by land–sea breeze circulations, shallow convection, and orographic flows (Jiménez et al. 2006). It may be expected that this layer is broken up by low-level convection and turbulence further downwind over Africa, mixing the pollution towards the surface, or transporting it further into the ITCZ (Kallos et al. 1998). The enhanced O3 levels recorded over the Mediterranean basin during MINOS have also been detected from space by the Global Ozone Monitoring Experiment (GOME) on the ERS-2 satellite. It appears that these remote sensing observations compare favourably with in situ measurements by aircraft and model results (Ladstätter-Weißenmayer et al. 2003). The coincident GOME and aircraft measurements of the O3 precursor gases NO2 and HCHO furthermore confirm the largescale importance of photochemical O3 formation from pollution sources (Kormann et al. 2003). These measurements point to very active photochemistry in the Mediterranean basin associated with strong oxidant production. It thus appears that O3 levels in the basin exceed the European Union eight-hourly air quality limit of 110 Ïg/m3 throughout most of the summer—and this is caused by European air pollution. In several parts of the basin, for example in the north-east of Spain, the EU plant protection threshold (80 Ïg/m) is exceeded more that 80 per cent of the time (Ribas and Peñuelas 2004). It is very difficult to control the ozone air quality locally, especially in the densely populated areas along the coast, since urban emissions add to the already high background O3 mixing ratios. In fact, based on scenario simulations for the year 2025, as shown in the right hand panel of Figure 22.5, it may be expected that tropospheric O3 will further increase, and that the Mediterranean region will remain as one of the ozone ‘hot spots’ in the Northern Hemisphere (Lelieveld and Dentener 2000). In these scenario calculations for 2025 it has been assumed that European NOx emissions will remain fairly constant, based on the expectation that technological improvements (such as limits on traffic NOx emissions), are compensated by an increasing

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number of vehicles and a growing fossil energy consumption. Remarkably, even though the European NOx emissions may not increase in the future, the ozone concentrations over the Mediterranean may nevertheless

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Fig. 22.5. Model-calculated ozone at the surface, averaged over the period May to August, for the present and the possible future (2025) atmosphere (after Lelieveld and Dentener, 2000).

increase as a consequence of enhanced background levels, caused by the rapidly growing ozone formation in Asian pollution air being transported on a hemispheric scale.

Air Pollution and Climate

High Atmospheric Oxidation Capacity Ozone, being a greenhouse gas and a prime oxidant, is also the main precursor of other oxidants such as the hydroxyl radical (OH), playing a key role in the photochemical degradation of species such as CO, hydrocarbons, and various sulphur compounds. Radical reaction chains can be initiated by the photo-dissociation of O3 , followed by primary OH formation from the reaction of excited oxygen atoms with water vapour (O1 D + H2 O → 2OH). The reaction chains are propagated through the catalytic action of nitrogen oxides (NOx ), which recycles the OH radicals (through HO2 + NO → OH + NO2 ) and additionally produces O3 by the photo-dissociation of NO2 . Since the primary formation of OH is controlled by ultraviolet radiation, ozone, and water vapour (all of which are abundant over the Mediterranean Sea in summer), and because OH recycling by NOx is also relatively efficient, the OH concentrations are quite high. Berresheim et al. (2003) measured an average OH concentration of 4.5(±1.1) × 106 molecules per cm3 during the MINOS campaign on Crete, with a strong diurnal cycle and high daytime peak values, as shown in Figure 22.6. This mean OH concentration is about a factor of three higher than the typical values for this latitude in summer. The associated concentration of sulphuric acid (H2 SO4 ), formed from SO2 + OH, is also quite high—up to 9 × 107 molecules per cm3 and this is

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among the highest reported in the literature (Bardouki et al. 2003). Note that H2 SO4 is the prime precursor of sulphate aerosol particles, discussed in the next section. During the day oxidation processes are dominated by OH radicals. However, during the night the OH concentration is very low so that atmospheric chemical pathways depend on other oxidants such as nitrate radicals (NO3 ) formed from the reaction between NO2 and O 3 . Sunlight photodissociates the NO3 radicals within seconds, so that its concentration follows an opposite diurnal cycle from that of OH (Figure 22.6). The night-time mean NO3 concentration during MINOS was 1.1(±1.1) × 108 molecules per cm3 . This is more than an order of magnitude higher than the daytime mean OH concentration of 8.2(±1.6) × 106 molecules cm3 (Vrekoussis et al. 2004). Since the reaction rate coefficients of non-methane hydrocarbons with NO3 are between 5 and 1,000 times slower than with OH, some species are preferentially converted by NO3 radicals. Vrekousis et al. (2004) calculated that, diurnally averaged, the conversion of dimethyl sulphide (DMS) by NO3 is only 33 per cent slower than that by OH, and that the further reaction of NO3 into N2 O5 and its heterogeneous conversion into HNO3 accounts for more than 20 per cent of the total nitrate production. It is interesting to see in Figure 22.6 that the variability of OH during the day is substantially less than the variability of NO3 during the night. This is because OH concentrations are ‘buffered’ through a balance between primary formation (O1 D + H2 O) and recycling through

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the catalytic action of NOx , even though the former process is dominant in the Mediterranean boundary layer (Lelieveld et al. 2002b). In summary, the oxidation capacity of the Mediterranean atmosphere is relatively high, during both day and night. The coincidence of reactive carbon compounds and nitrogen oxides, and the favourable conditions for oxidant formation promote the photochemical transformation of primary air pollutants such as NMVOC and SO2 . Although the lifetime of primary pollutant gases is thus limited, it leads to the build-up of reaction products such as ozone and aerosol precursors, which are inefficiently removed owing to the relatively slow dry deposition over the sea and the scarcity of rain. Therefore, the Mediterranean basin in summer, with its local pollution (re)circulation systems in a stable stratified atmosphere, is a ‘cooking vessel’ of photochemical smog.

Extensive Aerosol Haze Based on satellite measurements, the Mediterranean Sea has been identified as one of the maritime regions in the world with the highest aerosol optical depths (Husar et al. 1997). The earliest atmospheric chemistry measurements on Crete indicated relatively high concentrations of sulphate and nitrate, attributed to long-distance transport of air pollution (Mihalopoulos

et al. 1997). Upwind in northern Greece high levels of SO2 have been observed, attributed to coal burning in Central and Eastern Europe (Zerefos et al. 2000; Formenti et al. 2001). A small though significant fraction of 5–25 per cent of the sulphate, however, originates from natural DMS emissions by marine phytoplankton (Ganor et al. 2000; Kouvarakis and Mihalopoulos 2002; Chapter 2). Furthermore, desert dust intrusions from Africa, and additional—though smaller—fluxes of mineral dust from the Near and Middle East substantially contribute to the aerosol column (Dayan et al. 1991; Formenti et al. 2001; Chapter 14). Over Africa, the dust is often lifted to altitudes well above the Mediterranean boundary layer in synoptic disturbances. Hence, after its northerly transport from Africa it can entrain into the westerly flow over the Mediterranean in the free troposphere. It appears that the interannual variability of aerosol optical depth over the Mediterranean is strongly affected by the atmospheric column abundance of mineral dust, which correlates with the phase of the North Atlantic Oscillation that influences the atmospheric transport regime (Moulin et al. 1997). During the MINOS campaign, the fine aerosol mass consisted of more than one-third of sulphate and nearly one-third of particulate organic matter (POM), and it included substantial fractions of ammonium, black carbon, and other compounds (Figure 22.7). The sulphate was not fully neutralized and mostly present as

Fig. 22.7. Mean particle composition during the MINOS campaign for the fine (D < 2 Ïm) and the coarse (D ≥ 2 Ïm) mode aerosol based on measurements in Finokalia, Crete. The fine aerosol mass is 16.6 Ïg per m3 and the coarse aerosol mass is 24.2 Ïg per m3 (after Lelieveld et al. 2002a). NIS represents non-identified species.

Air Pollution and Climate

ammonium bisulphate—as also observed by Kouvarakis et al. (2002b). The coarse aerosol fraction mostly consisted of mineral dust and sea salt, including significant fractions of nitrate and sulphate. Lelieveld et al. (2002a) estimated that the fine aerosol particle mass was about 80–90 per cent anthropogenic, whereas about 60–80 per cent of the coarse mode particle mass was of natural origin. The pollutant aerosol sources include urban, industrial, agricultural, and forest fire emissions. Fire maps from the MODIS satellite instrument (Moderate Resolution Imaging Spectroradiometer) show that, during MINOS, west and north of the Black Sea, in Bulgaria, Romania, and in the Ukraine and Russia, respectively, extensive biomass burning occurred, 2,000 km or more upwind of Crete. These emissions reached the measurement site at Finokalia in about four days (Sciare et al. 2003a). Aircraft and mountain-based (1500 m a.s.l.) measurements corroborated the presence of pollutants in a reservoir layer above the boundary layer, in which the aerosol particles originated from Western, Central, and Eastern Europe. In general, agricultural burning and forest fires—mostly anthropogenically ignited—in the Mediterranean region contribute strongly to the aerosol burden, especially during dry spells (Sciare et al. 2003b; Chapter 19). Black carbon can originate both from biomass and fossil fuel combustion. Sciare et al. (2003a) estimated that during MINOS more than half the black carbon originated from fossil fuel use and the remainder from biomass burning. The relatively large mass of POM, on the other hand, included many oxygenated hydrocarbons, indicative of aged photochemically processed air pollution and biomass burning aerosol (Schneider et al. 2004). Only a small part of the POM appears to be related to natural emissions and this is discernible in the particles as formate, acetate, and ≥ c9 compounds, which were present only in very small concentrations. The mixture of natural and human-made particles forms an extensive haze over the Mediterranean Sea that can clearly be discerned on satellite images (Figure 22.2). Often visibility is limited, even on distant islands. Note that Crete represents the southernmost part of Europe in the Mediterranean region, remote from pollution emissions on the continent, whereas the observed local aerosol concentration is close to the European Union air quality standard for particulate matter (PM10 ) of 55 Ïg/m3 . This concentration is not allowed to be exceeded for more than thirty-five days per year. The high concentrations in Crete may be considered indicative for the even stronger violations of the air quality standards further upwind.

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By performing altitude resolved radiometric measurements, Markowicz et al. (2002) inferred that the diurnally averaged shortwave radiative forcing at the surface by aerosol particles is −17.9 ± 2.1 W/m2 (Figure 22.8). This is mainly due to solar radiation scattering and absorption by particles in the fine mode aerosol, substantially reducing the surface heating. The radiative forcing at the surface compares to a forcing of −6.6 ± 2.1 W/m2 at the top of the atmosphere (TOA); the latter representing the overall loss of solar energy to space through aerosol radiation backscattering. The difference between the TOA and the surface forcing corresponds to an atmospheric heating of 11.3 ± 3.8 W/m2 , caused by the absorption of solar radiation, in which black carbon plays a key role (Sateesh and Ramanathan 2000). Note that the above-mentioned surface forcing of nearly 18 W/m2 during MINOS is somewhat less than that derived from satellite measurements over a thirtyyear period (25 W/m2 ), with a relatively larger forcing during summer than during winter (Tragou and Lascaratos 2003). This could imply that the aerosol forcing may have decreased in the 1980s and 1990s, consistent with estimates of SO2 emissions in Europe as shown in the upper part of Figure 22.9. Figure 22.8 contrasts the negative aerosol forcing with the positive radiative forcing over the Mediterranean exerted by greenhouse gases. Obviously, these strong radiative

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perturbations may be expected to influence the Mediterranean atmospheric heating profile and the moisture budget through changes in evaporation and cloudiness.

Air Quality, Climate, and the Water Cycle The solar radiation energy input into the Mediterranean Sea, as reduced by aerosol particles, constitutes a major driving force of the oceanic thermohaline circulation (Tragou and Lascaratos 2003; Chapter 2). Indeed, substantial changes in Mediterranean deep water formation have been observed, possibly affecting salinity transport to the North Atlantic Ocean through the Strait of Gibraltar (Lascaratos et al. 1999). It should be further investigated to what extent these changes have been caused by surface radiative forcings of aerosol pollution and regional rainfall anomalies and how they link to global climate change, for example, through the North Atlantic Oscillation (NAO). The Mediterranean basin is characterized by large climate gradients, and there is concern that climate change

will be associated with the intensification of extreme weather conditions (Chapter 18). Regional European climate scenario calculations for the twenty-first century indicate a relatively strong warming as compared to the rest of Europe (Schär et al. 2004), and the number of very hot days, in particular, may increase (de Castro et al. 2004; Chapter 3). Air quality is strongly correlated with the mean temperature through the occurrence of hot and stagnant anticyclonic conditions (Lin et al. 2001) and these are expected to increase in future (Stott et al. 2004). In addition, the hemispheric background level of air pollution may increase through, for example, long-distance transport of emissions from Asia (Lelieveld and Dentener 2000). Furthermore, since a positive phase of the NAO correlates with increased transatlantic air pollution transport, a possible future change in the NAO may have consequences for air quality both in North America and Europe (Li et al. 2002). Climate scenario calculations also indicate that precipitation may increase in the western and northern parts of the basin, more often released in torrential rain events (Chapter 18), while in the drier southern and eastern parts precipitation is expected to decrease even further (Millán et al. 2005). However, the models used did not account for the effects of aerosol particles on the energy and moisture budgets. As indicated above, the strong aerosol scattering and absorption in the Mediterranean basin reduces the surface heating and thus the sea surface temperature (SST). The lower part of Figure 22.9 shows that the long-term SST variability, as influenced by the NAO, was fairly regular in the period 1930–70. Subsequently a strong cooling phase occurred, which correlates with SO2 emissions in Europe and presumably also with the sulphate aerosol burden over the Mediterranean Sea. Since the solar energy absorbed by the sea is largely returned to the atmosphere through evaporation, the negative radiative forcing at the surface, caused by sulphates and other particulate matter, suppresses evaporation and atmospheric moisture transport. To assess the possible consequences of this aerosol effect on the regional water cycle, Lelieveld et al. (2002b) performed a sensitivity study by prescribing observed low and high SSTs as boundary conditions to a general circulation model. The results demonstrate that the Mediterranean SST is a sensitive influence on the amount of precipitation received downwind in, for example, the Middle East and the eastern Sahel zone (Figure 22.10). First, the negative SST anomalies appear to correspond to drought periods in North Africa in the 1970s (Long et al. 2000). Second, the positive SST anomalies in the 1990s correlate with a recovery from

Air Pollution and Climate

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these spells in the same period (Nicholson et al. 2000) and these are coincident with decreasing SO2 emissions and sulphate concentrations (Figure 22.9). Further studies will be needed to substantiate these links between aerosol pollution, SST anomalies, and perturbations of the water cycle. The same is true for several additional though poorly quantified effects that aerosol particles can have on clouds and climate (Ramanathan et al. 2001). For example, the absorption of solar radiation by black carbon, which heats and thus stabilizes the aerosol pollution layer, could lead to the evaporation of clouds (Vogelman et al. 2001; Koren et al. 2004). Indirect aerosol effects on clouds also include the precipitation efficiency; for example, a high particle abundance may inhibit rainfall or suppress ‘warm’ rain formation in convective clouds (Rosenfeld and Woodley 2000; Rosenfeld et al. 2001). The latter effect can extend the vertical development of deep convective clouds, which promotes ice and hail formation and lightning so that some of these clouds may invigorate into heavy thunderstorms that produce torrential rain (Andreae et al. 2004; Koren et al. 2005; Chapter 18). To what degree these interactions between air pollution, clouds, and climate are relevant for the Mediterranean basin needs to be determined through coordinated research programmes. Global, regional, and local aspects influence both air pollution and climate, and mitigation or adaptation strategies should be based upon integrated problem assessments that also account for land use and soil hydrology changes. The largest risk lies in the possibility that some of these aspects combine

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into destabilizing (positive) feedback mechanisms in the Mediterranean with potentially large consequences for a region that has been shown to be vulnerable to changes in air quality, climate, and the water cycle.

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H. (2003), Formaldehyde over the eastern Mediterranean during MINOS: Comparison of airborne in situ measurements with 3D-model results. Atmospheric Chemistry and Physics 3: 851–61. Kouvarakis, G. and Mihalopoulos, N. (2002), Seasonal variation of dimethylsulfide in the gas phase and of methanesulfonate and non-sea-salt sulfate in the aerosol phase measured in the eastern Mediterranean atmosphere. Atmospheric Environment 36: 929–938. Tsigaridis, K., Kanakidou, M., and Mihalopoulos, N. (2000), Temporal variations of surface regional background ozone over Crete island in the South-East Mediterranean. Journal of Geophysical Research 105: 4399–407. Vrekoussis, M., Mihalopoulos, N., Kourtidis, K., Rappenglueck, B., Gerasopoulos, E., and Zerefos, C. (2002a), Spatial and temporal variability of tropospheric ozone (O3 ) in the boundary layer above the Aegean Sea (eastern Mediterranean). Journal of Geophysical Research 107: 8137, doi: 10.1029/2000JD000081. Doukelis, Y., Mihalopoulos, N., Rapsomanikis, S., Sciare, J., and Blumthaler, M. (2002b), Chemical, physical and optical characterization of aerosol during the PAUR II experiment. Journal of Geophysical Research 107: 8141, doi: 10.1029/2000JD000291. Ladstätter-Weißenmayer, A., Heland, J., Kormann, R., von Kuhlmann, R., Lawrence, M. G., Meyer-Arnek, J., Richter, A., Wittrock, F., Ziereis H., and Burrows, J. P. (2003), Transport and tropospheric build-up of trace gases during the MINOS campaign: comparisons of GOME, in situ aircraft measurements and MATCH-MPIC-data. Atmospheric Chemistry and Physics 3: 1887–902. Lascaratos, A., Roether, W., Nittis, K., and Klein, B. (1999), Recent changes in deep water formation and spreading in the eastern Mediterranean Sea. Progress in Oceanography 44: 5–36. Lawrence, M. G., Rasch, P. J., von Kuhlmann, R., Williams, J., Fischer, H., de Reus, M., Lelieveld, J., Crutzen, P. J., Schultz, M., Stier, P., Huntrieser, H., Heland, J., Stohl, A., Forster, C., Elbern, H., Jakobs, H., and Dickerson, R. R. (2003), Global chemical weather forecasts for field campaign planning: predictions and observations of large- scale features during MINOS, CONTRACE, and INDOEX. Atmospheric Chemistry and Physics 3: 267–89. Lelieveld, J. and Dentener, F. J. (2000), What controls tropospheric ozone? Journal of Geophysical Research 105: 3531–51. and 30 others (2002a), Global air pollution crossroads over the Mediterranean. Science 298: 794–9. Peters, W., Dentener, F. J., and Krol, M. C. (2002b), Stability of tropospheric hydroxyl chemistry. Journal of Geophysical Research 107: 4715, doi: 10.1029/2002JD002272. Li, Q. B., Jacob, D. J., Bey, I., Palmer, P. I., Duncan, B. N., Field, B. D., Martin, R. V., Fiore, A. M., Yantosca, R. M., Parrish, D. D., Simmonds, P. G., and Oltmans, S. J. (2002), Transatlantic transport of pollution and its effects on surface ozone in Europe and North America. Journal of Geophysical Research 107, doi: 10.1029/2001JD001422. Lin, C.-Y., Jacob, D. J., and Fiore, A. M. (2001), Trends in exceedances of the ozone air quality standard in the continental United States. Atmospheric Environment 35: 3217–28. Long, M., Entekhabi, D., and Nicholson, S. E. (2000), Interannual variability in rainfall, water vapour flux, and vertical motion over West Africa. Journal of Climate 13: 3827–41.

Air Pollution and Climate Lövblad, G., Tarrasón, L., Tørseth, K., and Dutchak, S. (2004), EMEP Assessment Part 1: European perspective. Norwegian Meteorological Institute, Oslo (available from ). Luria, M., Peleg, M., Sharf, G., Siman Tov-Alper, D., Schpitz, N., Ben Ami, Y., Gawi, Z., Lifschitz, B., Yitzchaki, A., and Seter, I. (1996), Atmospheric sulfur over the East Mediterranean region. Journal of Geophysical Research 101: 25917–30. Markowicz, K. M., Flatau, P. J., Ramana, M. V., Crutzen, P. J., and Ramanathan, V. (2002), Absorbing Mediterranean aerosols lead to a large reduction in the solar radiation at the surface Geophysical Research Letters 29/20: 1968, doi: 10.1029/2002GL015767. Mihalopoulos, N., Stephanou, E., Kanakidou, M., Pilitsidis, S., and Bousquet, P. (1997), Tropospheric aerosol ionic composition above the Eastern Mediterranean Area. Tellus 49B: 314–26. Millán, M. M., Salvador, R., Mantilla, E., and Kallos, G. (1997), Photo-oxidant dynamics in the western Mediterranean in summer: Results from European research projects. Journal of Geophysical Research 102: 8811–23. Mantilla, E., Salvador, R., Carratalá, R., Sanz, M. J., Alonso, L., Gangioti, G., and Navazo, M. (2000), Ozone cycles in the western Mediterranean basin: Interpretation of monitoring data in complex coastal terrain. Journal of Applied Meteorology 39: 487– 508. Estrela, M. J., Sanz, M. J., Mantilla, E., Martin, M., Pastor, F., Salvador, R., Vallejo, R., Alonso, L., Gangioti, G., Ilardia, J. L., Navazo, M., Albizuri, A., Artiñano, B., Ciccioli, P., Kallos, G., Carvalho, R. A., Andrés, D., Hoff, A., Werhahn, J., Seufert G., and Versino, B. (2005), Climate feedbacks and desertification: The Mediterranean model. Journal of Climate 18: 684–701. Moulin, C., Lambert, C. E., Dulac, F., and Dayan, U. (1997), Control of atmospheric export of dust from North Africa by the North Atlantic Oscillation. Nature 387: 691–4. Mylona, S. (1996), Sulphur dioxide emissions in Europe 1880– 1991. Tellus 48B: 662–89. Nicholson, S., Some, B., and Kone, J. (2000), An analysis of recent rainfall conditions in ‘West Africa, including the rainy seasons of the 1997 El Niño and the 1998 La Niña years. Journal of Climate 13 2628–40. Nolle, M., Ellul, R., Heinrich, G., and Güsten, H. (2002), A longterm study of background ozone concentrations in the central Mediterranean—diurnal and seasonal variations on the island of Gozo. Atmospheric Environment 36: 1391–402. Ramanathan, V., Crutzen, P. J., Kiehl, J. T., and Rosenfeld, D. (2001), Atmosphere, aerosols, climate and the hydrological cycle. Science 294: 2119–24. Ribas, A. and Peñuelas, J. (2004), Temporal patterns of surface ozone levels in different habitats of the North Western Mediterranean basin. Atmospheric Environment 38: 985–92. Roelofs, G.-J., Scheeren, H. A., Heland, J., Ziereis, H., and Lelieveld, J. (2003), A model study of ozone in the eastern Mediterranean free troposphere during MINOS (August 2001). Atmospheric Chemistry and Physics 3: 1199–210. Rosenfeld, D. and Woodley, W. L. (2000), Deep convective clouds with sustained supercooled liquid water down to –37.5◦ C. Nature 405: 440–2. Rudich, Y., and Lahav, R. (2001), Desert dust suppressing precipitation: A possible desertification feedback loop. Proceedings of the National Academy of Sciences 98: 5975–80. Sateesh, S. K. and Ramanathan, V. (2000), Large differences in tropical aerosol forcing at the top of the atmosphere and earth’s surface. Nature 405: 60–3.

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Vrekoussis, M., Kanakidou, M., Mihalopoulos, N., Crutzen, P. J., Lelieveld, J., Perner, D., Berresheim, H., and Baboukas, E. (2004), Role of the NO3 radicals in oxidation processes in the eastern Mediterranean troposphere during the MINOS campaign. Atmospheric Chemistry and Physics 4: 169–82.

Ward, M. (1992), Provisionally corrected surface wind data, worldwide ocean-atmosphere surface fields and Sahelian rainfall variability. Journal of Climate 5: 454–75. Zerefos, C., Ganev, K., Kourtidis, K., Tzortsiou, M., Vasaras, A., and Syrakov, E. (2000), On the origin of SO2 above Northern Greece. Geophysical Research Letters 27: 365–8.

This chapter should be cited as follows Lelieveld, J. (2009) Air pollution and climate, in J. C. Woodward (ed.), The Physical Geography of the Mediterranean. Oxford University Press, Oxford, 599–614.

23

Biodiversity and Conservation Jacques Blondel and Frédéric Médail

Introduction The biodiversity of Mediterranean-climate ecosystems is of particular interest and concern, not only because all five of these regions (the Mediterranean basin, California, central Chile, Cape Province of South Africa, western and southern parts of Australia) are among the thirty-four hotspots of species diversity in the world (Mittermeier et al. 2004), but they are also hotspots of human population density and growth (Cincotta and Engelman 2000). This relationship is not surprising because there is often a correlation between the biodiversity of natural systems and the abundance of people (Araùjo 2003; Médail and Diadema 2006) and this, inevitably, raises conservation problems. Within the larger hotspot of the Mediterranean basin as a whole, ten regional hotspots have been identified (Figure 23.1). They cover about 22 per cent of the basin’s total area and harbour about 44 per cent of Mediterranean endemic plant species (Médail and Quézel 1997, 1999), as well as a large number of rare and endemic animals (Blondel and Aronson 1999). A key feature of these Mediterranean hotspots as a whole is their extraordinarily high topographic diversity with many mountainous and insular areas. Not surprisingly this results in high endemism rates and they contain more than 10 per cent of the total plant richness (see the recent synthesis of Thompson 2005). However, of all the mediterranean-type regions in the world, the Mediterranean basin harbours the lowest percentage (c.5%) of natural vegetation considered to be in ‘pristine condition’ (Médail and Myers 2004; Chapter 7). With an average of as many as 111 people per km2 , one may expect a significant decline in biological diversity in the Mediterranean basin—a region that has been managed, modified, and, in places, heavily degraded by humans for millennia (Thirgood 1981; Braudel

1986; McNeill 1992; Blondel and Aronson 1999; Chapter 9). There are two contrasting theories that consider the relationships between humans and ecosystems in the Mediterranean (Blondel 2006, 2008). The first one is the ‘Ruined Landscape or Lost Eden’ theory, first advocated by painters, poets, and historians in the sixteenth and seventeenth centuries, and later by a large number of ecologists. This theory argues that human action in the form of deforestation and overgrazing resulted in a progressive and cumulative degradation and desertification of Mediterranean landscapes. This theory denounces the destruction of the formerly extensive forests which were supposedly so lush and large that a monkey could have travelled from Spain to Turkey almost without leaving the crown of the trees! One example of this view has been vividly depicted by David Attenborough in his book, The First Eden: The Mediterranean World and Man (Attenborough 1987; see also Naveh and Dan 1973; Thirgood 1981; McNeill 1992). Challenging this pessimistic view, the second school dismisses the supposedly detrimental effects of humans by arguing that the imaginary past idealized by artists and scientists does not acknowledge the fact that humans actually contributed to the maintenance of Mediterranean landscapes as they were progressively established since the last glacial episode, stressing that savanna-like landscapes are characteristic of the Mediterranean (Grove and Rackham 2001). Reality is inevitably somewhere between these two extremes which are of crucial relevance to any discussion of threats to biodiversity in this region. In fact, several modelling attempts suggest that the Mediterranean region is one of the most threatened regions in the world with regards to environmental changes (Sala et al. 2000). Generally speaking, all the components of

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Maritime and Ligurian Alps

Western Mediterranean Islands Betic–Rifan complex High and Middle Atlas

Djurdjura– Kabylies

Southern and Central Greece Crete

Southern Anatolia and Cyprus

Syria, Lebanon and Israel

Mediterranean Cyrenaica

0

500 km

Fig. 23.1. The ten regional species diversity hotspots of the Mediterranean basin based on plant endemism and richness (based on Médail and Quézel 1997, 1999 and Véla and Benhouhou 2007).

‘global change’ (Vitousek 1994) threaten, to various extents, the biodiversity in this region, but the component of most concern is human population growth and its consequences for species, habitats, and ecosystems. Indeed, the resident population has increased from 213 million people in 1950 to 427 million in 2000, with 523 million expected in 2025 and 600 million in 2050, mainly located along the coasts (Benoit and Comeau 2005). Rather than reviewing all the threats that currently affect Mediterranean biotas, this chapter will focus on some of the most important drivers—such as habitat destruction and fragmentation, and the impacts of fire and invasive species—and try to make some generalizations about habitat dynamics and the impact of these threats upon the most important and/or wellknown groups of terrestrial and freshwater plants and animals.

Contrasted Human Impact in Space and Time What characterizes ecosystems and habitats much more in the Mediterranean region than in any other region in the world is their long-lasting common history with

humans as they have been designed and redesigned by them for almost 10,000 years in the eastern part of the basin and around 8,000 years in its western part (e.g. Le Houérou 1981; Braudel 1986; Pons and Quézel 1985). However, the action of humans has been far from always detrimental and has sometimes resulted in an increase in biodiversity through the shaping of a large variety of cultural landscapes. For example, Blondel and Aronson (1995, 1999) argue that many traditional land use practices act as surrogates of natural disturbance regimes with the consequence that, according to the intermediate disturbance hypothesis (Huston 1994), several components of biodiversity have actually been much higher in landscapes shaped by humans than in primitive plant communities such as oak woodlands. The coexistence of stock-farming, crop fields, and wildlife has led to higher local diversity in many parts of the Mediterranean so that abandoning sustainable and moderate land use practices inevitably decreases the biodiversity of many groups. Examples of traditional agrosylvo-pastoral land use systems include the Roman triad known as sylva-saltus-ager (woodland-pasture-field) in which landscapes are divided into distinct plots devoted to wood cutting (sylva), livestock breeding (saltus), or crop harvesting (ager). Another traditional system is the

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Fig. 23.2. A montado in Portugal with cattle under the trees and charcoal burners (photo: Jacques Blondel).

Montado-dehesa system of the Iberian Peninsula and several Mediterranean islands where the three activities mentioned above occur in the same plots (Figure 23.2). Thanks to the fine balance of woodlots, pastoral grasslands, shrublands, and open spaces reserved for cultivation, the resulting ‘moving mosaic’ greatly contributed to the biological diversity of Mediterranean landscapes (Blondel and Aronson 1995, 1999; Grove and Rackham 2001). Although many of these ancient land use practices have almost disappeared in most parts of the basin, their imprint on the structure of landscapes and habitats is still visible today. This is an illustration of the value for biodiversity of cultural landscapes and habitats and the evolving history of land use impacts by humans (Foster 2002). Over past millennia and increasingly as cultural impacts became more varied, ecosystems and habitats have changed continuously, supporting mosaics of highly dynamic landscapes (Blondel and Aronson 1999; Foster et al. 1990). Most of them have been influenced and shaped by cultural activity for so long that it is

often difficult to decipher the natural from the human components of their history. The impact of human land use and disturbance legacies is apparent in all ecosystems (Foster 2002). For example, in an archaeological study of French forests, Dupouey et al. (2002) demonstrated that the signature of agriculture and pastoralism dating back to the Romans is still visible today in the form of a mosaic-like distribution of certain plant species (e.g. Vinca minor, Ribes uva-crispa) which are typically linked to ancient human settlements. Furthermore, some recent palaeoecological studies suggest that the forest cover was not as dense and uniform as formerly thought during the Holocene optimum, with many landscapes being rather open and heterogeneous, but also with some strong regional differences in vegetation dynamics (Beaulieu et al. 2005). However, the idyllic vision of a ‘natural sustainability’ driven by traditional agro-pastoral systems is not easy to sustain as shown by the many examples of ancient and devastating human impacts on Mediterranean ecosystems (Thirgood 1981; Chapter 9).

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The present trends of habitat changes as a result of human impact vary markedly across the regions of the Mediterranean basin (Barbero et al. 1990; Mazzoleni et al. 2004; Blondel 2008). In the northern parts, the collapse of traditional land use systems and rural depopulation from the end of the nineteenth century, and especially after the two world wars, completely destructured traditional landscapes as they were replaced by industrial-style agriculture and modern activities such as mass tourism. Since the beginning of the twentieth century, changes in land use practices in the European parts of the Mediterranean included industrial-scale cultivation, but also a rapid transformation of previously cultivated or grazed lands towards shrubby and forested areas, leading to what may be described as a homogenization of the flora and fauna. As a result, thermophilous taxa of Mediterranean origin tend to be replaced by more ubiquitous species of medioEuropean origin. One example is the replacement of Mediterranean warblers—which are typical of Mediterranean shrubby habitats—such as the Sardinian warbler Sylvia melanocephala or the Spectacled warbler Sylvia conspicillata by forest species such as the Blackbird Turdus merula or the Blue tit Cyanistes caeruleus (Blondel and Farré 1988; Debussche et al. 2001). In the northern, European side of the Mediterranean, human pressure on many ecosystems has steadily decreased as a result of agricultural abandonment and rural depopulation dating back to the end of the nineteenth century, which has accelerated greatly since the Second World War. Across the entire range of life zones and habitat types, a progressive recovery of forests and matorrals invading old fields, and abandoned terraces and vineyards is evident (Mazzoleni et al. 2004; Figure 23.3; Chapter 7). These processes lead to a decrease in habitat patchiness and a slowing down of the ‘moving mosaic’, that is the turnover in time of habitat patches according to how farmers manage them, which is so characteristic of Mediterranean landscapes and beneficial for several components of biological diversity. The collapse of traditional land use systems that maintained habitat heterogeneity during centuries undoubtedly had and will have consequences on biodiversity and its distribution at the landscape scale. In contrast, in North Africa and most of the eastern part of the Mediterranean, pressures on natural habitats by human populations and livestock are still strongly on the increase, destructuring soils and ecosystems and resulting in intense erosion. Woodward (1995) mentioned that the impact of ongoing woodland clearance followed by overgrazing has resulted in high rates of soil erosion across the Mediterranean drainage basin.

The average area annually cultivated with cereals in North Africa increased by 50 per cent during 1948– 52 and 1981–90 (Le Houérou 1990). This resulted in more extensive clearing of arid steppe rangeland which is a major cause of desertification in North Africa and the Near East. As a result, disturbance regimes in farm, pasture, and forest lands are moving towards still greater intensity of land and resource exploitation for the shortterm survival of local people. With an average population growth of 3.2 per cent per year in North Africa and several countries of the Near East, the consequences of increased human pressure in the arid zone of the southern shores of the Mediterranean could be extensive runoff, frequent flooding, and extreme degradation of the remaining patches of forest (Figure 23.4). Soil erosion, which averages 5–10 t ha−1 yr−1 in medium to large catchments, might increase by a factor of five or more to reach average values of 25–50 t ha−1 yr−1 as measured in limited areas of Algeria and Morocco (ibid.) where 35 per cent of land is already undergoing losses exceeding 30 t ha−1 yr−1 (see Chapter 20). These opposite trends that characterize the dynamics of ecosystems on the two sides of the Mediterranean have markedly contrasting impacts on the wealth of biological systems with a trend of habitat regeneration in the north and a trend of ongoing degradation in the south. In the northern part of the basin, forest is recovering at a rate of 2 per cent per year as a result of rural depopulation and land abandonment, whereas it is almost exactly the opposite in the eastern and southern parts of the basin (Quézel and Médail 2003).

Habitat Destruction All habitats and landscapes in the Mediterranean region, except perhaps some remote mountainous and steep cliff areas, have been to some extent managed and transformed by humans. These tremendous human-induced changes have had profound consequences on biodiversity. Although some of them have had beneficial consequences on the dynamics and distribution of species, communities, and habitats as mentioned above, many have resulted in serious threats to biodiversity. Forests, wetlands, and coastal habitats, together with their associated flora and fauna, are perhaps those that have been, and continue to be, most affected by deforestation, land reclamation, and human encroachment on coastlines (Ramade 1997). A gap analysis suggests than no more than 5–10 per cent at most of the native postglacial Mediterranean forest is left (WWF 2001) with huge geographical

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Fig. 23.3. Forest recovery on ancient terraces (photo: Jacques Blondel).

variation. The main threats to forests are mismanagement associated with overexploitation of resources such as timber, firewood, and grazing—especially in the southern and eastern parts of the basin (Quézel and Médail 2003). A particularly dramatic situation is

that of the forests or forest-steppes of the Maghreb with native woodland dominated by Cedrus atlantica, Quercus canariensis, Quercus faginea, or Juniperus thurifera being reduced to half of their native extent or seriously disturbed, and the former beautiful forests of Tetraclinis

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Fig. 23.4. The last individual of a former forest of Juniperus thurifera in southern Morocco (photo: Frédéric Médail).

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articulata have almost disappeared (Quézel 2000). The dramatic decline of the famous cork oak forest of the Mamora in northern Morocco in only seventy-five years is a striking example of the magnitude of forest destruction in North Africa. This forest extended over 130,000 ha in 1920 but was reduced to 77,064 ha in 1960, 60,000 ha in 1980, and a mere 12,844 ha in 1995. The degradation and draining of wetlands is not a recent problem but has a history going back several centuries during which many coastal wetlands have disappeared (Chapter 9). For example, the Etruscans began to drain the marshes around Rome as early as the fifth century AD (Pearce and Crivelli 1994). Most of the 250 freshwater or brackish lakes and marshes of Central Anatolia (180 of which are more than 100 ha in size) have already disappeared, many of them having been drained in the 1950s and 1960s in order to tackle malaria and to develop new agricultural land. Between 1984 and 2002, the surface area of the Anatolian marshes of the Konya plain was reduced by 90 per cent and the most striking case is that of the 10,000 ha of the Yarma marshes which were completely dried up within ten years (Gramond 2002). Many wetlands that served as refugia during glacial times for Euro-Siberian and boreal plant species (e.g. Andrieu-Ponel et al. 2000) and which were initially designated by Roi (1937) as colonies planitiaires (e.g. the lower Rhône in southern France, the Arno and Serchio estuaries in Tuscany, and the Pontin marshes near Rome) have been drained, sending many local populations of relict flora and fauna to extinction. In the early 1920s, Greek wetlands covered three times their present area, and a third of the remaining wetlands are threatened. In Macedonia alone, 1,151 km2 of wetlands out of a total of 1,572 km2 have been drained since 1930 (Catsadorakis 2003). Because wetlands are very productive ecosystems, they are also much desired by humans and many of them are managed for the production of various crops. For example, the delta of the River Axios in Greece is increasingly threatened by the expansion of the large city of Thessaloniki with mussel beds and ricefields providing 90 and 70 per cent respectively of the total Greek production. In spite of these encroachments, the delta is still an important breeding and stopover place for 250 species of birds, 76 of which are rare, and 36 species of fish. Even in prestigious and well-protected areas such as the Camargue, habitat losses have not stopped since the nineteenth century with a net loss of c.40,000 ha during the second half of the twentieth century (Tamisier and Grillas 1994; Figure 23.5).

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The high consumption of water for agriculture and tourism induces an over-exploitation of groundwater around the Mediterranean, especially in Israel, Spain, Malta, Cyprus, Tunisia, and southern Italy, contributing to extensive drainage of wetlands. Water consumption peaks during the summer, when the resources are at their lowest, with each Mediterranean tourist using between 300 and 850 litres of water per day (De Stefano 2004). The result of this tremendous increase in water demand has been the destruction of 1 million hectares of wetlands within the Mediterranean region over the last fifty years (Skinner and Zalewski 1995; Chapter 21). Riparian forests are associated with wetlands and they share a great many plant and animal species. These environments are well known for their diversity as well as for their role in controlling floods. In addition, they act as natural corridors allowing many species of temperate Europe to colonize Mediterranean habitats, greatly contributing to their biological diversity (Décamps and Décamps 2002). Unfortunately most of the Mediterranean riparian forests have been destroyed and replaced by intensive agriculture or by stands of trees such as poplars. Among other aquatic habitats of particular interest in the Mediterranean are the temporary and oligotrophic ponds, which constitute species-rich and original habitats with several rare and endangered ferns (Pilularia, Isoëtes, Marsilea), phanerogams (Damasonium), insects, frogs, and large branchiopods such as Triops cancriformis (Grillas et al. 2004; Figure 23.6). They are also key breeding sites for amphibians, including the rare Mediterranean newt Triturus marmoratus (Figure 23.7). Their characteristic species are adapted to the drastic and unpredictable conditions imposed by summer aridity, interannual fluctuations in hydrology, and a poor supply of nutrients. These temporary marshes are indeed of major conservation importance, but they have been considerably degraded by various disturbance events, especially in North Africa (Rhazi et al. 2001; Grillas et al. 2004). Such environments will also be threatened by climate change given the scenarios associated with a warmer climate in the region (Chapter 3). Coastal habitats and ecosystems are also increasingly under threat. Rapid changes in land use practices in the twentieth century, especially over the last four decades, have had disastrous consequences for coastal ecosystems where more than 60 per cent of people live. In south-eastern France, for example, population density reaches 2,500 people per km2 in the Alpes-Maritimes département compared with an average of just 108 people per km2 in the country as a whole (Figure 23.8).

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Fig. 23.5. Wetlands are among the most threatened habitats in the Mediterranean region (photo: Jacques Blondel).

In Italy, 70 per cent of the coast is already urbanized and rates of human encroachment on coastal areas are growing at an unprecedented rate in many Mediterranean regions such as southern Spain, the coasts of the Adriatic Sea, and many countries of the Near East. Rampant urban development and pollution are serious threats to coastal habitats because in many areas urban waste is pumped untreated directly into the sea. Forming a zone where terrestrial and aquatic habitats meet, the coast is especially fragile and vulnerable. For example, 60 per cent of the Greek population, 40 per cent of Greek agriculture, 70–80 per cent of Greek industrial activity, and 90 per cent of tourism concentrate in the coastal zone (Catsadorakis 2003). The accelerated rate of urbanization of coastlines across the Mediterranean is likely to reach 75–80 per cent almost everywhere in the basin by 2025. The high demand for Mediterranean coastal and insular landscapes—because of their beauty and popularity with tourists—makes them particularly threatened (Delanoë et al. 1996), especially because of their well-known vulnerability owing to a high pro-

portion of rare endemic species and unique insular habitats.

Habitat Fragmentation Landscape fragmentation is a very ancient practice everywhere around the Mediterranean, adding to the natural diversity of habitats, especially along the coasts and around urban areas. If fragmentation may threaten biodiversity by increasing inbreeding and extinction risks for small isolated populations (Noss and Csuti 1994), the process is not necessarily detrimental overall, as shown by the consequences of traditional land use practices on several components of biodiversity such as alpha, beta, and gamma diversity (see Blondel and Aronson 1995, 1999). For example, the fragmentation of oak forests in southern France resulted in a local increase of species richness of many groups, especially ruderal plant species that benefit from edge effects along roads and crop fields where nitrogen-rich nutrients are imported by humans. However, the richness of different plant functional types in the core areas of forests

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Fig. 23.6. The crustacean branchiopod Triops cancriformis with the rare thrumwort Damasonium stellatum and the parsley frog Pelodytes punctatus in a temporary marsh in the Camargue (Blondel and Aronson 1999).

remains higher or equal to that of the edges (Médail et al. 1998). Fragmentation may seriously change the dynamics of biological interactions. For example, as a result of high densities of seed eaters (‘supersaturation’) such as voles in habitat mosaics resulting from fragmentation, seed recruitment of forest species may be modified. In Mediterranean Spain, the average number of seedlings of the holm oak Quercus ilex is nine times higher in relatively large undisturbed forests (150–350 ha) than in small forest fragments (0.2–12 ha in size) where most acorns are eaten by animals (Santos and Tellería 1997). Post-dispersal loss may also be exacerbated within small isolated habitats. For example, up to 25 per cent of the fruits of Juniperus thurifera may be eaten by voles (Apodemus sylvaticus) in small habitat fragments, whereas such losses do not exceed 5 per cent in large forest tracts (Santos and Tellería 1994).

Habitat fragmentation may also have serious consequences on the fitness, genetic diversity, and local adaptation of organisms. For example, in Cyclamen balearicum, a rare endemic plant species from southern France and the Balearic Islands, isolation and fragmentation of sclerophyllous oakwoods resulted in a loss of genetic diversity, an increase of genetic drift, and selection-biased mechanisms leading to spontaneous self-pollination in highly isolated populations (Affre et al. 1997; Thompson 2005). A long-term programme has investigated the causes and consequences of habitat fragmentation on phenotypic plasticity (that is, the production by the same genotype of different phenotypes depending on local habitat features) in a song bird, the Blue tit (Cyanistes caeruleus). Results showed that depending on the size of patches within habitat mosaics and the type of oak, whether deciduous (Quercus humilis) or evergreen (Q. ilex), this

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Fig. 23.7. A rare amphibian of temporary ponds of the western part of the basin, the Mediterranean newt, Triturus marmoratus (photo: N. Zbinden).

small passerine responds in several ways to habitat fragmentation. In a mainland landscape in southern France, where birds are assumed to disperse freely across habitat patches, little local differentiation of two breeding traits—laying date and clutch size—has been found within a range of c.40 km, suggesting gene swamping between populations (Blondel et al. 2001; Figure 23.9). A molecular genetics study of populations within this landscape supported the hypothesis of a source-sink population structure with more birds immigrating from deciduous habitat patches to evergreen ones than the reverse (Dias et al. 1996). In a similar geographic configuration of habitats in Corsica, there was a much higher phenotypic variation and a higher degree of population differentiation on a scale that is usually smaller than the dispersal range of Blue tits. This difference between mainland France and Corsica has been interpreted as resulting from reduced dispersal ranges of birds on islands and supports the divergence-with-gene-flow model of speciation (Blondel et al. 1999; Blondel 2008). In the mainland landscape where mismatching between breeding time and the peak of food abundance is high, maladaptation can result in poor breeding success making such populations dependent on source populations at the scale of landscapes. This implies that conserva-

tion strategies must consider the geographical configuration of habitats at large spatial scales. The responses of Blue tits to habitat patchiness show how organisms can respond to environmental changes including climate change. The changes that happened many times during the Pleistocene are currently occurring at an accelerated speed as a result of global warming. This may create problems for organisms across the Mediterranean region as they attempt to adapt (Jump and Peñuelas 2005). All the empirical evidence suggests that the biological consequences of habitat fragmentation should be carefully studied, because they are complex and generally underestimated.

Is Fire a Real Threat to Mediterranean Biodiversity? On average 1 per cent of the Mediterranean forests, that is c.600,000 ha, burn annually with many consequences for human societies. Wildfires frighten people and have detrimental socio-economic effects, but provided their frequency does not exceed a certain threshold they are major components of the dynamics of Mediterranean-type ecosystems and do not necessarily

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Fig. 23.8. The urbanization of coastal areas threatens habitats and rare plants and animals almost everywhere in the Mediterranean basin. On this photo the city of Monaco is built on a habitat for the rare woody shrub Euphorbia dendroides that forms the bright rounded clumps in the foreground (photo: Frédéric Médail).

threaten biodiversity (Moreno and Oechel 1994; Eugenio and Lloret 2006; Chapter 19). The use of fire by humans is a very old practice in the Mediterranean region, perhaps as ancient as 700,000 BP in northern Israel (Goren-Inbar et al. 2004). However, regular use of fire became more frequent only between c.7,000 and 5,000 BP, mainly to clear vegetation for agricultural and pastoral purposes, but also as a strategic weapon in the wars of Antiquity (KuhnholtzLordat 1938; Pons and Thinon 1987). In modern times, c.50,000 wildfires occur annually, most of them caused by humans—only 5 per cent at most are natural fires generated by lightning. The majority of burnt hectarage is caused by a few very large wildfires: in coastal eastern Spain, 144 large fires (i.e. only 1.3% of the total) destroyed 487,793 ha (i.e. 78% of the total burnt surface areas) between 1968 and 1994 (Piñol et al. 1998; Chapter 19). From a landscape perspective, Mediterranean fires are disturbance events that maintain habitat heterogeneity

and the moving mosaic that keeps ecosystems functioning. For example, Naveh (1999) demonstrated, in his studies in the Mount Carmel area of Israel, the key role of fire as an evolutionary and ecological factor in shaping landscapes and vegetation. Many communities and species of open habitats, including xero-thermophilous species, are narrowly dependent on such disturbance events. One example of the role of fire in maintaining bird communities in Mediterranean landscapes has been provided by Prodon et al. (1987). At the scale of the whole range of habitats within a landscape consisting of a habitat mosaic generated by recurrent fire events in southern France, fifty-one species of birds have been found, each of them occupying its preferred habitat. Bird species of open vegetation such as larks (e.g. Lullula arborea), partridges (e.g. Alectoris rufa), and linnets (Carduelis cannabina) colonize habitats immediately after fire and are then replaced by species of matorrals such as warblers (Sylvia spp.) and the Nightingale (Luscinia megarhynchos) as habitats recover. In turn, matorral

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Laying date (1=1st of March)

100

80

CE1

CE3 Evergreen

CD1 ME1

CE2

60

ME2 CD2

40

MD1

ME3

MD2 Deciduous

20 Corsica

Mainland

Fig. 23.9. Mean laying date (in ‘March-date’, i.e. 32 = 1 April) of the Blue tit Cyanistes caeruleus in mainland and Corsican habitats. The horizontal dotted lines indicate the best date for onset of breeding relative to resource availability (leaf-eating caterpillars) in deciduous habitats (D) and in evergreen habitats (E), respectively. CE = ‘Corsica evergreen’, CD = ‘Corsica deciduous’, ME = ‘Mainland evergreen’, MD = ‘Mainland deciduous’. Vertical bars = 1 standard deviation (modified from Blondel et al. 2001).

species will be replaced by species of later transitional stages and so forth until the recovery process is completed, returning to a forest stage with such species as nuthatches and woodpeckers. Fire events periodically move up and down the position of any given habitat patch on the grassland → matorral → forest system. At a broader geographical scale, the combination of a disturbance regime and community dynamics, resilience, and inertia results in an equilibrium that may be fairly stable in the long term (Blondel 1987). Communities of small mammals have been shown to recover quickly after fire with a definite sequence of local colonizations and extinctions that match successional changes in the structure of vegetation (Prodon et al. 1987; Torre and Diaz 2004). Several species of small mammals, especially the wood mouse Apodemus sylvaticus and to a lesser extent the white-toothed shrew Crocidura russula and the Algerian mouse Mus spretus, have been shown to be more abundant in the early post-fire habitats than later on in transitional stages of the post-fire habitat gradient, presumably as a result of the combination of large quantities of seeds and seedlings in these herbaceous areas, as well as alleviated predation pressures in recently burnt areas (Torre and Diaz 2004). An amazing and unexpected fact in burnt habitats is the local persistence after fire of Mediterranean land snail communities (Kiss et al. 2004; Kiss and Magnin 2006), presumably because of the availability of underground microrefuges.

The post-fire resilience of Mediterranean vegetation is a well-known characteristic that is due to several intrinsic properties of plants (Trabaud and Prodon 2002; Eugenio and Lloret 2006). These include the resprouting capacities of plants with the existence of lignotuber, the role of the edaphic seed-bank, and the colonization of burnt areas by seed dispersal through wind or vertebrates (Quézel and Médail 2003). Only very high fire return rates can induce a real malfunctioning of ecological processes resulting in threats to biodiversity and increased runoff and soil erosion. Repeated fires combined with soil erosion have caused drastic edaphic changes in some postglacial soils and have contributed to a strong decline in Mediterranean old-growth forests formed in part by laurifolious (e.g. Ilex aquifolium, Taxus baccata, Laurus nobilis) and deciduous (e.g. Quercus spp., Celtis australis, Cercis siliquastrum) trees. One of the consequences of erosion and ecosystem degradation on limestone soils is the so-called leopard’s skin effect, with an alternation of bare ground and stones with low bushes of dwarf oaks (Quercus coccifera) or spiny xerophytes (e.g. Astragalus, Sarcopoterium, Ulex) as shown in Figure 23.10. To the authors’ knowledge, however, no plant species has become extinct as a result of fire. In fact, some rare and threatened plants are actually favoured by fire events. Examples in southern France include several species of Fabaceae (e.g. Genista linifolia, Vicia altissima, Vicia melanops) whose seed germination is enhanced by high temperatures, which is also the case for most of the Mediterranean geophytes (e.g. Acis, Iris, Orchis, Serapias, Tulipa). A detailed study of a rare endemic geophyte of the Maritime Alps (Acis nicaeensis, Amaryllidaceae) indicated that, a year after the event, the fire induces a significant increase in the density of flowering individuals and seedling emergence, as well as in clump densities, by reducing aboveground plant competition and increasing bare soil cover (Diadema et al. 2007). However, such differences were quickly attenuated two years after the fire event. These results suggest that small-scale fires can be beneficial for the regeneration window of this threatened geophyte, through the periodic supply of the seed and bulb banks in the soil.

Invasive Species Of all the threats to biodiversity in the Mediterranean, the introduction of alien species as a side-effect of human colonization has been particularly harmful for many biotas. The introduction of exotic species can cause considerable disruption to native ecosystems (e.g. Drake

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Fig. 23.10. Intense habitat degradation due to repeated fire events results in very low scrubby vegetation and bare ground, the so-called leopard’s skin (photo: Jacques Blondel).

et al. 1989; Williamson 1996) so that biological invasions are currently considered as one of the main threats to biodiversity (Mooney and Hobbs 2000). In some regions of the world, especially on islands, invading species have displaced whole cohorts of native species and disrupted ecosystem dynamics; thus they are increasingly considered potential threats to biodiversity and ecosystem functioning. However, this assertion was recently questioned by Gurevitch and Padilla (2004) who argue that on a worldwide scale endangered taxa are more than five times more threatened by habitat loss than by biological invasions, with alien species contributing less than 2 per cent to the extinction of the 762 documented plants and animals in the world IUCN (International Union for Conservation of Nature) database. Thus, the primary role of invasive aliens in driving widespread extinctions has not yet been clearly demonstrated; more research is needed to estimate the real threats more accurately. On the whole, the Mediterranean basin appears to be less vulnerable to species invasion than the other

mediterranean-type regions of the world, especially California and the Cape Province of South Africa. It has been suggested that the low invasiveness of mediterranean ecosystems, except for some habitats such as freshwaters, wetlands, or riparian forests, results from the strong interactions that tied humans and ecosystems for millennia, with recurrent human-induced disturbance regimes, resulting in some kind of coevolution between them (di Castri et al. 1990). Combined with thousands of spontaneous colonization events and selection pressures by humans, disturbance regimes made ecosystems progressively more resistant with ‘old invaders’ preventing potential ‘new invaders’ from colonizing them (Drake et al. 1989). Among fish species, ancient introduced species such as the carp (Cyprinus carpio) or the perch (Perca fluviatilis) can be considered as naturalized, but many recently introduced species pose real threats for the native fauna. Invading taxa may account for more than half the species in some water bodies. For example, nineteen species in Spain as compared to thirty-two

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native, and ten species in Portugal as compared to twenty-nine native have been introduced (Crivelli and Maitland 1995). The situation seems to be better in some other regions such as Greece where only 11 of the 117 fish species of the country have been introduced. Some species are particularly harmful for the indigenous fish fauna, including predators such as the sheatfish (Silurus glanis) the pike-perch (Sander lucioperca), and the bass (Micropterus salmoides). The introduction of herbivorous carp from China in 1983 in Lake Oubeira, Algeria, resulted in the destruction of half the reed-beds of this marsh (Pearce and Crivelli 1994) with disastrous consequences for the native biotas. A severe drought completely desiccated this lake some years ago, however, killing all fish species including the introduced carp. The Louisiana crayfish (Procambarus clarkii) deserves special attention—it invaded many rivers and wetlands of the region, causing irreversible damage to many communities of invertebrates. Exotic crayfish species are an example of the paradox of invasions because they may benefit populations of emblematic and popular predators such as herons that feed on them. Unfortunately the negative side is an undetected and irreversible damage to other components of communities that make ecosystems function. Relatively few exotic birds and mammals have invaded Mediterranean ecosystems. Even the Pheasant (Phasianus colchicus), which was introduced by the Romans, did not succeed in establishing self-sustaining free populations and the many attempts to introduce species as game birds (the Californian quail Colinus virginianus, the Black francolin Francolinus francolinus, and some others) usually met with failure. The only species that established themselves in the wild are small populations of the Ring-necked parakeet Psittacula krameri in several large cities around the Mediterranean, free-ranging populations of the Red avadavar Amandava amandava in Spain, and the Ruddy duck, Oxyura jamaicensis, from North America. It can be argued that the last example should be eliminated because of risks of hybridization with the native endangered White-headed duck Oxyura leucocephala. Among mammals, some species have been successfully introduced among which the Coypu Myocastor coypus and the Musk rat Ondatra zibethicus may cause damage in wetlands. Attempts to introduce the Floridan rabbit Sylvilagus floridanus as a game species have not been very successful except in some scattered localities. With only 250 naturalized plants in ‘natural’ areas, i.e. 1 per cent of the total plant richness (Quézel et al. 1990), the Mediterranean region is generally considered less vulnerable to invasion by alien plant species

than the other four mediterranean-type regions of the world, but the problem may be slightly underestimated. It is also noteworthy that the flora of the Mediterranean basin often provides an important source of alien species to other mediterranean-type regions, particularly California. The global threats induced by alien plants to rare or endemic Mediterranean plants are still relatively small (Thompson 2005). Indeed, the well-documented survey performed within the project Atlas y libro rojo de la flora vascular amenazada de España (Bañares et al. 2003) mentions that only 8 per cent of the studied populations (n = 2,223) of rare plants are threatened by the competition induced by xenophytes. However, some exotic plant invaders are a serious threat to coastal and riparian ecosystems, with Mediterranean islands being more threatened by invading plants than most mainland areas (Lloret et al. 2004). For example, exotic plant species represent 17 per cent (473 taxa) of the Corsican flora (Natali and Jeanmonod 1996), but only 1.4 per cent (38 taxa) of them are considered as naturalized in the natural vegetation. Nevertheless, numerous alien plants are becoming invasive in some parts of the Mediterranean basin, inducing profound ecological disruptions in natural or semi-natural habitats (Table 23.1). Some South African Carpobrotus taxa (Aizoaceae) were largely introduced at the beginning of the nineteenth century, and are often naturalized, causing one of the most severe invasive threats in Mediterranean coastal and island habitats (Suehs et al. 2001). Studies of genetic variation suggest that C. acinaciformis individuals are of hybrid origin in some islands off the coast of France, and have generated recent introgressed types through repeated backcrossing with C. edulis genotypes (Figure 23.11). All these strains together form a large and variable hybrid swarm with an aggressive invasiveness (Suehs et al. 2004a, b). Thus, the rapid evolutionary potential of hybrid genotypes can constitute a serious additional threat by making hybrid plants particularly invasive. Wetlands are particularly sensitive to invasion and have been colonized by several xenophytes some of which cover huge areas, causing much damage to freshwater plant and animal communities. Examples are Baccharis halimifolia, Cortaderia selloana, Ludwigia grandiflora, and L. peploides, which can form thick carpets completely covering rivers and canals. Coming across an angler recently near a canal in the Camargue, one of us saw the man pulling a bass out of a thick carpet of Ludwigia. It appeared that the belly of the fish was full of Louisiania crayfish. The three players of this scene were three exotic species which presumably developed at the expense of many native species.

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TABLE 23.1. Some of the most invasive alien plants occurring in the Mediterranean basin Taxa (and main species) Acacia spp. (A. dealbata, A. karroo, A. longifolia, A. melanoxylon, A. saligna . . . ) Agave americana Ailanthus altissima Araujia sericifera Arundo donax Aster squamatus Azolla filiculoides Baccharis halimifolia Buddleja davidii Carpobrotus spp. (C. edulis, C. acinaciformis) Conyza spp. (C. bonariensis, C. canadensis, C. sumatrensis) Cortaderia selloana Cotula coronopifolia Datura stramonium Eucalyptus spp. (E. camaldulensis, E. globulus . . . ) Gomphocarpus fruticosus Ipomoea ssp. (I. sagittata, I stolonifera, I. purpurea, I. indica) Isatis tinctoria Lemna minuta Lonicera japonica Ludwigia spp. (L. grandiflora, L. peploides) Medicago arborea Nicotiana glauca Oenothera spp. (O. biennis, O. glazioviana, O. dillenii) Opuntia ssp. (O. ficus-indica, O. dillenii, O. stricta, O. subulata, O. crassa . . . ) Oxalis pes-caprae Paspalum spp. (P. dilatatum, P. paspalodes, P. vaginatum) Pennisetum villosum Phytolacca americana Pittosporum tobira Ricinus communis Robinia pseudoacacia Senecio angulatus Senecio inaequidens Sorghum halepense Solanum spp. (S. chenopodioides, S. bonariense, S. sisymbrifolium . . . ) Tradescantia fluminensis Xanthium spp. (X. italicum, X. spinosum)

Family

Native distribution

Main invaded habitats in the Mediterranean basin

Mimosaceae

Australia, S. Africa

Woodlands, matorrals, coastal habitats

Agavaceae Simaroubaceae Asclepiadaceae Poaceae Asteraceae Azollaceae Asteraceae Buddlejaceae Aizoaceae Asteraceae

C. America China S. America C. Asia S. and C. America S. and N. America N. America China S. Africa S. and N. America

Coastal habitats, matorrals Riparian or ruderal habitats, woodlands Matorrals, evergreen forests Ruderal and riparian habitats, agroecoystems, old-fields Ruderal habitats, wetlands, old-fields, sandy beaches Freshwater Wetlands, riparian habitats Riparian or ruderal habitats Coastal habitats Ruderal habitats, agroecoystems, old-fields

Poaceae Asteraceae Solanaceae Myrtaceae

S. America S. Africa America Australia

Ruderal habitats, wetlands, old-fields, matorrals Wetlands Ruderal and riparian habitats Woodlands, matorrals

Asclepiadaceae Convolvulaceae

S. Africa S. America

Riparian or ruderal habitats Coastal matorrals, wetlands

Brassicaceae Lemnaceae Caprifoliaceae Onagraceae Fabaceae Solanaceae Onagraceae

S.E. Asia N. and S. America E. Asia S. and N. America E. Mediterranean S. America S. America

Ruderal habitats, old-fields, agroecosystems, matorrals Freshwater Wetlands, riparian habitats Freshwater Coastal habitats, matorrals Coastal and ruderal habitats Ruderal or riparian or habitats, matorrals, old-fields

Cactaceae

C. and S. America

Rocks, matorrals, coastal habitats

Oxalidaceae Poaceae

S. Africa S. America

Agroecosystems, old-fields Agroecosystems, old-fields, wet pastures, wetlands

Poaceae Phytolaccaceae Pittosporaceae Euphorbiaceae Fabaceae Asteraceae Asteraceae Poaceae Solanaceae

N.E. Africa N. America E. Asia Palaeotropical N. America S. Africa S. Africa E. Mediterranean C. and S. America

Ruderal habitats, old-fields Riparian and ruderal habitats, mesophilous matorrals Coastal habitats, matorrals Ruderal or riparian habitats Woodlands, riparian habitats, matorrals Coastal matorrals, evergreen forests Ruderal habitats, agroecoystems, old-fields Ruderal habitats, agroecoystems, old-fields Ruderal or riparian habitats

Commelinaceae Asteraceae

S. America America

Riparian habitats, wetlands, rocks Ruderal or riparian habitats, sandy beaches

Source: Modified from Médail, unpublished data.

The Decline of Mediterranean Biodiversity Except for large mammals and some endemic species on Mediterranean islands, relatively few postglacial extinction events due to human impact have been reported in the Mediterranean in recent times (Chapter 5). This does not mean, however, that the status of Mediterranean

biodiversity is safe. Changes of anthropogenic origin have had many impacts on the distribution and abundance of populations, making a large number decline while others are increasing. In fact, for many plants, some groups of vertebrates, and most groups of invertebrates, micro-organisms and fungi, there is much uncertainty on their status and several supposedly extinct species may survive in unnoticed localities, or as

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Fig. 23.11. Carpobrotus acinaciformis is a very aggressive invasive plant species in coastal areas (photo: Frédéric Médail).

dormant seeds or bulbs in plants as reported for example by Greuter (1994) who mentioned that several plant species thought to have vanished have recently been rediscovered.

Plant Biodiversity Despite a large number of endemic plant species that are narrowly distributed in a single or few localities, few taxa seem to have become extinct in the Mediterranean region, except perhaps on some islands such as the Canary islands and Madeira, where 24 and 20 per cent respectively of native plant species are considered to be threatened. The first census performed by Greuter (1994) indicated than only thirty-seven Mediterranean plants were presumed totally extinct, i.e. an extinction rate of 0.13 per cent of the native Mediterranean flora (Table 23.2). The present survey, taking into account several recent floristic checklists and red data books, allowed us to identify forty-two species and subspecies that are probably extinct within the strict Mediterranean bioclimatic region, which represents an

extinction rate of 0.15 per cent. Sixteen of these taxa were not cited by Greuter (1994) and only half of them are common in the two surveys. The largest number of plant extinctions have been reported for Turkey (10 taxa), Greece, and Italy (6 taxa each). However, the number of documented cases of extinctions in countries of North Africa and the Middle East is presumably disproportionately low given the ongoing huge loss of habitats. The current level of extinction is perhaps not significantly higher than during earlier geological periods (Chapter 4), but it is very difficult to estimate the magnitude of postglacial extinctions for plants. Among the very few cases of local extinctions during historical time, the disappearance of Tetraclinis articulata stands from Cyrenaica (Libya) was already testified by Theophraste in c.500 BC. This tree of the Cupressaceae family occurs today only in the Maghreb—its native status is doubtful in Andalucia. The strong local persistence and resilience of Mediterranean plant species is mainly due to their life-strategies, with a high tolerance of stress and disturbance (Chapter 7). One example is the

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TABLE 23.2. Threatened vascular plants by country based on the former IUCN categories and included in the 1997 IUCN Red List of Threatened Plants Country

Ext

Ex/E

E

Albania Algeria Croatia Cyprus Egypt France Gibraltar Greece Israel Italy Jordan Lebanon Libya Malta Morocco Portugal Spain Syria Tunisia Turkey

1

3

1 31

2 7 6

3 1 2

1 1 1 1 2 3 2 10

3

1

9 19 21 28 7 29 1 2 4 3 46 185 5 1 47

V

22 2 13 8 81 3 80 7 80 1 18 1 3 113 272

167

R

I

73 80 4 23 45 83 1 430 13 190 6 1 30 10 157 98 484 3 18 161

2 8

Total threatened plants

Total plant species

79 141 6 51 82 195 4 571 32 311 9 5 57 15 186 269 985 8 24 1, 876

6 10 7 32 3 12 1 2 5 4 23 12 41 5 53

% threatened

3, 031 3, 164 5, 347 1, 682 2, 076 4, 630 — 4, 992 2, 317 5, 599 2, 100 3, 000 1, 825 914 3, 675 5, 050 5, 050 3, 000 2, 196 8, 650

2.6 4.5 0.1 3 3.9 4.2 — 11.4 1.4 5.6 0.4 0.2 3.1 1.6 5.1 5.3 19.5 0.3 1.1 21.7

Note: The figures also include the plant species that are present in the non-Mediterranean areas of these countries. Source: Modified from Walter and Gillett 1998.

amazing persistence of only c.200 individuals of Zelkova sicula, a recently discovered Tertiary relict tree located in only one small area in south-east Sicily (Garfi 1997) (Figure 23.12). This extreme case of persistence through time of a plant species results from its vegetative propagation, which combines a clonal structure and a process of adaptive reiteration with a continuous replacement of the vegetative apex which is repeatedly destroyed by drought or grazing. Even if it is very difficult to have a precise idea of the number of plants that are threatened in the Mediterranean, some details are available for certain areas or plant categories. In the southern shore countries, which are still in a process of ongoing degradation as explained above, Leon et al. (1985) estimated that nearly 25 per cent of the Mediterranean flora may be threatened in the decades to come, but precise and updated appraisals are still lacking. The unique flora of Mediterranean islands is on the whole threatened, especially the endemic species that grow on coastal and low-altitude habitats: on large islands, the percentage of taxa that are threatened on a global scale ranges from 2 per cent (Corsica) to 11 per cent (Crete) (Table 23.3). A detailed comparison of life-form spectra for the floras of south-east France and Corsica in relation to altitudinal distribution and rarity (Verlaque et al. 2001), indicates that endemic plants

are not more prone to extinction than other species. Extinction rates and rarity percentages of endemics in Corsica are lower than those of non-endemic taxa (1.6 v. 3.8% and 33 v. 37.5% respectively). Rare plants (i.e. occurring in less than ten localities) are mainly located at low altitude, a sensitive zone with the highest extinction rates. As much as 87 per cent of the extinct plants in Corsica and 90 per cent in south-east France occurred between 0 and 800 m above sea level, mainly in arable fields, wetlands, coastal areas, and rocky grasslands.

TABLE 23.3. Threatened vascular plants on the seven large Mediterranean islands, based on the former IUCN categories Island

Balearic Corsica Sardinia Sicily Crete Malta Cyprus

Ext

E

V

R

I

Total

Threatened taxa/total flora (%)

1 1 0 1 0 1 0

10 8 11 11 11 0 9

14 27 30 26 61 1 14

43 10 21 45 118 10 22

1 1 1 4 3 4 6

69 47 63 87 193 16 51

5 2 3 3 11 2 3

Note: Ext: extinct; E: endangered; V: vulnerable; R: rare; I: insufficiently documented. Source: Delanoë et al. 1996.

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Fig. 23.12. Ramet of Zelkova sicula (Ulmaceae), a relict and very threatened palaeoendemic small tree located in south-eastern Sicily (from Raimondo et al. 1994 in Quézel and Médail 2003).

The global extinction rate of plant species is higher in south-east France (4.79%, n = 161 local extinct plants) than in the less disturbed island of Corsica (3.48%, n = 74 local extinct plants). This is probably due to the more ancient and important human impact on continental areas, which include 7.3 times more inhabitants. On the whole, plants without visible vegetative organs during unfavourable seasons (annuals, aquatic, and bulbous groups) are the most threatened, with a rate of population extinction reaching 21 per cent for aquatic plants in south-east France (ibid.). Considering Mediterranean trees, the World List of Threatened Trees (Oldfield et al. 1998) includes thirtyfive Mediterranean species, and the 1997 IUCN Red Data Book mentions twenty-nine tree species (Walter and Gillett 1998), but a more precise list includes sixty-one taxa with forty-two endemics (Quézel and Médail 2003),

some of them being progenitors of cultivated trees (Malus, Olea, Phoenix, Prunus, Pyrus). One of the most noteworthy of these is Phoenix theophrasti, an eastern Mediterranean palm recognized by Theophrastus and Pliny but only formally described in 1967 by Greuter. According to ancient texts, the former distribution of this palm was more extensive, but nowadays this relict tree persists only in scattered populations at low altitude (0–350 m a.s.l.), in riparian habitats, and moist cliffs within some coastal sites in Crete, the Peloponnese, and south-west Turkey (Boydak 1987). Several populations are threatened by the continuous increase of tourist pressure along the coasts of Turkey and Greece, especially water harnessing, and by the putative increase of drought events associated with global warming (Chapter 3). Threats upon the emblematic cedar species are also evident (Quézel and Médail 2003). The famous Lebanon cedars were exploited for several millennia by the Egyptians, Phoenicians, Assyrians, Romans, and Turks who used large volumes of timber for ship-building and other uses. Only fourteen isolated and fragmented stands of Lebanon cedar currently remain, covering at most an extremely reduced surface area of 2,700 ha. The conservation of bulbous plants is another concern around the Mediterranean. In the Near East, notably in Turkey, intensive harvesting of geophytes for horticultural or culinary purposes still constitutes a serious problem since many of these taxa are endemic. In the 1980s in Turkey alone, 60 to 80 million bulbs (mainly Galanthus, Eranthis, Leucojum, Anemone, and Cyclamen) were collected in the wild each year, and 57 million tubers belonging to thirty-eight species of orchids were picked annually for the preparation of salep, a popular milk drink (Sezik 1989). Some integrated conservation programmes including in situ bulb propagation and then resale by villagers have reduced by half these huge removals and constitute an alternative source of bulbs for the international trade (Entwistle et al. 2002). The narrow distribution of many Mediterranean geophytes increases their extinction risks, even if storage organs such as bulbs or rhizomes may enhance population persistence over time. For example, the western Mediterranean snowflakes (Acis, Amaryllidaceae) encompass nine bulbous plant species and almost all of these are endemics that are threatened by various factors, including severe human impacts such as overgrazing and deforestation. Acis tingitana, which is a north Moroccan endemic, is especially vulnerable to such impacts. Other kinds of land use change that encourage rapid habitat colonization by competitive herbs and shrubs are serious threats for Acis nicaeensis and Acis fabrei (Figure 23.13),

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agricultural landscapes, and the putative extinction of several endemic or relict plants, with several rare native cytotypes. On the whole, however, data allowing a predictive overview of plant extinctions in relation to patterns of rarity and threats in the Mediterranean are still lacking. Future work should be inspired by the detailed study on Centaurea corymbosa, a narrow endemic and cliff-dwelling Asteraceae from southern France (Colas et al. 1997, 2001). The low seed dispersal and reduced ability to establish new populations due to obligatory cross-pollination (allogamy) and a single reproductive event in the life of the plant (semelparity) represent a serious handicap, enlarged by an Allee effect (negative effect of low population density on demographic parameters) on seed production. As the metapopulation dynamics of this species is very slow and depends upon only six natural populations located within an area of 3 km2 , the survival of this species is indeed in the balance.

Animal Biodiversity

Fig. 23.13. Acis fabrei, a narrow endemic with only four known populations from the southern slopes of the Mont Ventoux (Vaucluse, south-eastern France) (photo: Frédéric Médail).

which are narrow endemics of southern France, and Acis longifolia, which is an endemic from northern Corsica. Mediterranean segetals are plant species, mostly annuals, that are associated with cereal crop fields. They encompass over 1,500 species but their decline is also worrying as a result of land use changes that have induced profound shifts in the floristic composition of arable fields, olive groves, and other traditional agricultural lands. Changes to cultural practices can favour the spread of invasive weeds, mainly polyploid plants characterized by a high competitive capacity, at the expense of the traditional segetals. The latter are often stress-tolerant and annual species, autogamous, diploid, and with an autumnal germination (e.g. Camelina, Cephalaria, Delphinium, Garidella, Hypecoum, Nigella, Tulipa). The disappearance of this highly specialized flora contributes to the decline of Mediterranean

Historical records, both palaeontological and archaeological, show that a large human-induced decline in animal biodiversity started many thousands of years ago in most parts of the Mediterranean. The end of glacial times was characterized in the Northern Hemisphere by the presence of an extraordinary large number of large mammals including no less than twenty-five herbivorous as well as predatory species (Table 23.4). Most of them are now extinct but a testimony of this magnificent Late Pleistocene mammal fauna is provided by dozens of fossil sites in southern Europe as well as by the superb wall paintings in many ornate caves of southern France and northern Spain. Two caves with beautiful paintings have recently been discovered in southern France, the Chauvet cave in 1994 and the Cosquer cave in 1996. The Chauvet cave contains some of the oldest (about 30,000 BP) and most beautiful cave paintings ever found in the Mediterranean basin (Chauvet et al. 1995). A vast bestiary is portrayed, giving a vivid picture of what the fauna of this epoch looked like, including three hundred or more different animals such as bisons, horses, bears, deer, mammoths, hyenas, panthers, lions, rhinos, reindeers, aurochs, and ibexes, as well as the only known representation in Palaeolithic art of large birds such as the Eagle owl. In the underwater Cosquer cave near Marseilles, paintings depict many mammals such as horses, bisons, and ibexes, and an extraordinarily detailed rendition of the Great auk. This painting, which dates back 20,000 years, testifies that breeding colonies of this species, which is now extinct, occurred on the

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Jacques Blondel and Frédéric Médail TABLE 23.4. List of large mammals that were present in the Mediterranean basin during the Late Pleistocene, including species found as fossils in various deposits of southern France and which became extinct (E) before the Holocene Families and species Canidae Canis lupus Vulpes vulpes Alopex lagopus Cuon alpinus Ursidae Ursus thibetanus Ursus arctos Ursus spelaeus Hyaenidae Crocuta spelaea Felidae Panthera (Leo) spelaea Panthera pardus Lynx spelaea Felis silvestris Mustelidae Gulo spelaeus Meles meles Proboscidae Palaeoloxodon antiquus Mammuthus primigenius Rhinocerotidae Coelodonta antiquitatis Dicerorhinus hemitoechus

Late Pleistocene

Extinct

+ + + +

+ + E E

+ + +

E + E

+

E

+ + + +

E + E +

+ +

E +

+ +

E E

+ +

E E

Families and species Equidae Equus sp Equus germanicus Suidae Sus scrofa Cervidae Rangifer tarandus Megaceros giganteus Cervus elaphus Capreolus capreolus Bovidae Bos primigenius Bison priscus Hemitragus sp. Hemitragus cedrensis Capra sp. Capra ibex Rupicapra rupicapra Rupicapra sp. Sciuridae Marmotta marmotta Castor fiber

Late Pleistocene

Extinct

+ +

+ E

+

+

+ + + +

E E + +

+ + + + +

E E E E +

+ +

+ E

+ +

+ +

Source: Modified from Defleur et al.1994 in Blondel and Aronson,1999.

shores of the Mediterranean Sea during glacial times. Some survivors of this ancient fauna, which included a surprisingly high number of large predators, disappeared only recently, since the lion and elephant survived until historic times in Greece and Syria, respectively. Perhaps the most dramatic and still controversial event of the Late Pleistocene is the tempo and mode of the mass extinction of a large part of the rich megafauna that was so characteristic of this epoch (Chapter 5). The question whether this mass extinction was mainly caused by humans as advocated by the overkill hypothesis (Martin 1984) or resulted from dramatic environmental changes, or a combination of both, is still open. As pointed out by Martin and Klein (1984), large body size repeatedly appears as a characteristic associated with higher rates of extinction during episodes of environmental change. Extinction events in Europe and North America during the Late Pleistocene differentially eliminated mammals of large body size and open steppe and grassland. Birds that disappeared at the same time also included a disproportionate number of carnivores and scavengers as well as species of large size and open habitats. The decimation of large endemic mammals continued well into the Holocene, however, as sadly

illustrated by the human-induced extinction of all the large ‘mega-nano-mammals’ of Mediterranean islands, for example the dwarf hippos and elephants of Cyprus, Malta, Sicily, following the colonization of these islands by humans in the early Holocene, some 10,000 years ago (Simmons 1988, 1991). In addition to these mammals, tortoises and flightless owls which populated most Mediterranean islands were also decimated by humans as soon as they invaded the islands (Blondel and Vigne 1993; Vigne et al. 1997). In more recent times, the combination of habitat changes and direct persecution has doomed to extinction many species of large mammals, especially in North Africa. For example, the Brown bear Ursus arctos, the Ass Equus asinus, several species of gazelles (Gazella rufina, Oryx dammah, Alsephalus busephalus), the Lion Panthera leo, the Porcupine Hystrix cristata, the Cuvier’s gazelle Gazella gazella, and the Mouflon Ammotragus lervia have been extirpated from many parts of the Mediterranean, especially North Africa (see Blondel and Aronson 1999). Species that suffer directly from persecution are top predators such as the Panther Panthera pardus and the Monk seal Monachus monachus which are on the verge of extinction. The wolf, which is the

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Fig. 23.14. Three taxa of salmonids: (top) the Marble trout which is the most endangered fish species of the Adriatic basin, (middle) the Corsican trout with pure genotypes only in the upper reaches of river basins, and (bottom) the Brown trout which is still widespread across the region.

animal most closely associated with wilderness areas, has already disappeared from fourteen European countries and is in danger of disappearing from several others. Viable populations still persist in the Balkans and, with 500–700 individuals, Greece has the second largest population in the European Union after Spain. The species is currently reoccurring in France following immigration from Italy in the 1990s. Freshwater fish are perhaps the most interesting group of Mediterranean vertebrates, but also one of the less known and the most threatened with more than 300 species, 44 per cent of them being endemic (Crivelli and Maitland 1995). Several families include a large number of endemic species, for example the eighty-

three species of Cyprinidae, twelve species of Cobitidae, and eight species of Salmonidae (Crivelli 1996), with beautiful but endangered species such as the Marble trout Salmo marmoratus and the Corsican trout Salmo trutta macrostygma (Figure 23.14). Unfortunately, 70 per cent of these endemic species are seriously threatened, and four species are already extinct (ibid.). Most threats are from the degradation of water both in quality (pollution, eutrophication) and quantity (huge amounts of water are pumped for domestic use and irrigation), hybridization and invading species, especially predatory species. Other threats to freshwater animals are habitat destruction and loss combined with pollution, which are particularly harmful to the troglobiotic (cave-dwelling)

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fauna along Dalmatian coasts in countries bordering the Adriatic Sea, especially in Slovenia and Croatia. In limestone areas, cave waters as well as interstitial waters are contaminated by diffuse pollution from the use of fertilizers, pesticides, and, episodically, spills (Sket 1999). Commercial collecting of cave beetles and of one of the most fascinating Mediterranean amphibian, the blind cave salamander Proteus anguinus, can also be locally important threats. Annexes of the EU Habitat Directive listing species of ‘European Importance’ tend to ignore this kind of fauna, which is particularly diverse and specific to the karst areas of Mediterranean Europe (Chapter 10). Reptiles and amphibians have not suffered serious extinction events in recent times, since only the Crocodile Crocodilus niloticus and the endemic Blackbelly toad Discoglossus nigriventer have disappeared from Israel and southern Morocco, and marshy areas of the Israel–Syria border, respectively (Blondel and Aronson 1999). However, most populations of amphibians and many populations of reptiles are declining

everywhere in the world, including the Mediterranean (Balmford et al. 2003; Grillas et al. 2004). The main threats include epidemics, habitat acidification, desiccation, climate change, and the introduction of exotic species. For example, the massive introduction of Slider turtles Trachemys scripta elegans into Europe as pets often induces the release of these exotic turtles into natural habitats where they have been successful in reproducing. Competition with the European pond turtle Emys orbicularis results in weight loss and high mortality in the latter where the two species co-occur (Cadi and Joly 2004; Figure 23.15). Another threat for amphibians is the introduction of predatory fish species in amphibian breeding sites which are naturally fish-free. In such sites amphibians develop inducible defences that are costly to produce, potentially affecting the population dynamics of species (Teplitsky et al. 2003). Among birds, no more than three large species in North Africa (the Lappet-faced vulture Torgos tracheliotus, the Helmet guinea fowl Numida meleagris, and the Arabian bustard Choriotis arabs) have disappeared, while

Fig. 23.15. The pond terrapin Emys orbicularis, a vulnerable species in decline (photo: H. Hafner).

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Fig. 23.16. A pair of Bonelli’s eagles Hierraaetus fasciatus at their nest. This species is a typical cliff-dwelling raptor which is declining in most parts of its range (photo: Jacques Blondel).

a few others such as the Demoiselle crane Anthropoides virgo and the Bald ibis Geronticus eremita are in danger of extinction. There are, however, many changes in the structure and composition of the bird fauna of the Mediterranean basin. Many species that were formerly widespread everywhere in the basin are now confined to small, localized populations, many of them being threatened, mostly as a consequence of habitat loss or change, pollution, and sometimes food shortage. A serious decline in bird populations has been reported for many species and this decline may be catastrophic for some groups such as farmland birds (Donald et al. 2001). Some species are particularly vulnerable, for example the Bonelli’s eagle Hierraaetus fasciatus (Figure 23.16) and the Lesser kestrel Falco naumanni with a decline that has accounted for 95 per cent of the European population since 1950 (BirdLife International 2000). For large insectivorous birds such as the Lesser kestrel and several species of shrikes, the main cause of decline is food shortage resulting from the intensification of agriculture and urban sprawl (Liven-Schulman et al. 2004). For migrant birds, pesticide use in Africa to fight against

locusts might affect birds through a sharp reduction in prey availability. On the whole, there is a general trend of decline in population sizes, especially waders, sub-Saharan migrants, and farmland species. Despite some progress made in the European Union through the Birds Directive, the conservation status of many birds in Europe has worsened since 1994, and 43 per cent of Europe’s 526 species are now under threat (BirdLife International 2004), which is 5 per cent more than in 1994 (Tucker and Heath 1994). Socio-economic and political changes may also seriously affect some components of biodiversity as illustrated by the story of Griffon vultures and transhumance. The Griffon vulture Gyps fulvus used to belong to a kind of trophic community closely linked to migrating herds of wild mega-herbivores such as the European Bison, Auroch, Mouflon, Wild goats, Saiga antelope, and Tarpan. The up and down transhumance of these animals between lowlands in winter and highlands in summer occurred annually in April–May and October. The carcasses and dung of these animals— which were scattered along migration routes—provided

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plenty of food for vultures, corvids, mammals, insects (e.g. dung beetles), and other animals associated with these resources. Transhumance of domestic herds of sheep and cattle probably followed the same transhumance routes as wild herds and persisted in Mediterranean Europe up to the middle of the twentieth century, providing plenty of food for vultures and associated animals in many parts of the Mediterranean. For example, at least nine large colonies of the Griffon vulture have been reported to occur along the gorges of the Drina River in the Balkan Peninsula, as shown by many sub-fossil remains spreading from the upper Pleistocene and Holocene to modern times (Marinkovic and Karadzic, 1996). Political circumstances with changes in borders combined with shorter migration routes and sedentarization of cattle herds resulted in a decline and abandonment of transhumance. In the 1960s, laws forbidding transhumance in the former communist federal state of Yugoslavia further dramatically reduced the abundance of nomadic herds which subsequently completely disappeared. The break-up of the long migratory routes of cattle on the Balkan Peninsula started at the end of the nineteenth century after the collapse of the Ottoman Empire and the formation of new national borders. This, combined with a widespread use of poison (e.g. strychnine) against large carnivorous mammals, resulted in a complete collapse of the animal communities associated with these long-distance mammal migrants. On the whole, for birds as well as for many insects and reptiles, the main consequences of human-induced changes in Mediterranean habitats over time have not been so much a decrease in overall species richness at a regional scale than a tremendous advantage for species adapted to drylands and shrublands, as opposed to forest-dwellers. In some cases, direct persecution resulted in the death of millions of birds. Particularly sad examples are the large-scale destruction of migrant birds in Mediterranean islands such as Malta, Cyprus, and others where migrating birds, including the tiny kinglets, used to provide a providential amount of food for local people twice a year (Fenech 1993). Magnin (1991) has estimated that up to 1,000 million birds are killed annually in the basin when they stop over during migration. Except for very few exceptions such as butterflies and some large families or genera of conspicuous beetles, almost nothing is known on the distribution and abundance status of most groups of invertebrates. For groups that have been studied in some detail, there is a growing body of evidence for a severe decline (Collins and Thomas 1991; Pullin 1995). For example, a dramatic

decline has been reported for large conspicuous insects such as butterflies, large bees, dragonflies, and many groups of beetles, including all saproxylophagic groups that are narrowly tied to decaying wood (Speight 1989). The decline in population abundance of most groups of large insects paralleled that of large insectivorous birds such as the Roller Coracias garrulus, the Little owl Athene noctua, the Scops owl Otus scops (Figure 23.17), and all species of shrikes. Another dramatic decline of invertebrates is that of earthworms (Granval and Muys 1992) which has been attributed to leaching, acid rain, heavy metal pollution, and all factors associated with intensive mechanized agriculture. Given their key function in ecosystems, the decline of earthworms and the concomitant soil compaction contributes to an increase in soil erosion and severe flooding (Chapter 6).

Genetic Changes and the Decline of Human-selected Plant and Animal Varieties Subtle changes potentially detrimental to biodiversity may result from (1) introgression between native and cultivated populations as is the case with some northern populations of Oleander Nerium oleander, or (2) hybridization between closely related species that secondarily come in contact as a result of human-induced changes in the environment. For example, Mediterranean orchids (Ophrys, Orchis, Dactylorhiza, and Serapias) have a wide range of isolating barriers that lower the risks of interspecific recombination under normal circumstances. But in many regions of the basin, isolating barriers may have become ineffective because of millennia of human activities that have altered landscape structure and created disturbed sites conducive to hybridization. On the other hand, a recent aspect of biodiversity decline is that of the regression of species, cultivars, and various breeds of economic value that are now neglected after centuries of selection by humans. The Mediterranean basin is a key area for the presence of many wild progenitors of cultivated plants and for the origin and spread of agriculture (Zohary and Hopf 2000). In many species, adaptive intraspecific variation occurred as a response to human-induced selection and habitat changes over millennia, resulting in the differentiation of a burst of local ecotypes and gene pools in plant species, with region-specific characters fitting them to local climate and environmental conditions (Chapter 7). Vavilov (1935) enumerated a large number of plant taxa that differentiated as a result of human selection through

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Fig. 23.17. The Scops owl Otus scops is threatened by the decline of large invertebrates which constitute the bulk of its food (photo: Jacques Blondel).

the process of domestication. Over the centuries, hundreds of varieties of olive, almond, wheat, grape, etc., were passed down if not preserved in hortus. In many cases, the horticultural and agricultural selections and discoveries of past generations have been lost in recent decades. For example, only a very few of the 382 named cultivars of almond that were in use at the turn of the last century on the island of Mallorca are left (Socias y Company 1990). Throughout the Mediterranean area, the great majority of ‘minor’ fruits, nuts, vegetables, and other plant varieties selected in the past are extinct and

lost forever. The loss of these ancient varieties is an issue of great concern and much effort is devoted today to restore and preserve them from extinction. In the same way, another threat that directly concerns humans is the extinction of many livestock breeds or varieties. Of 3,381 breeds or varieties (horses, donkeys, cows, buffalo, pigs, goats, and sheep) recorded worldwide up to the twentieth century, 618 (16%) were already extinct—many of them in the Mediterranean, which is the region with the greatest variety of breeds. The Mediterranean basin includes 45 per

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cent of the bovid varieties and 55 per cent of the goat varieties of Europe and the Middle East but a large proportion of them are in danger of extinction (Georgoudis 1995), mainly because European dairy breeds dramatically outnumber of local breeds. Greece still has a total of thirty-nine livestock breeds, for example, but has already lost eight in the last century (Catsadorakis, 2003).

Conservation Framework and Programmes for Conserving Biodiversity and Restoring Endangered Species and Populations Many protected areas have been created around the Mediterranean basin, with various levels of protection, and there are currently seventy-six National Parks and forty Man and Biosphere Reserves (Table 23.5). Much effort has also been done in recent years to protect wetlands in the framework of the Ramsar convention, and there are at present 206 Ramsar sites representing a total area of more than 4 million hectares (Table 23.6; Chapter 9). Protected forest areas amount to c.10 per cent of forest coverage in the north shore countries of the MediterTABLE 23.5. Major protected areas such as National Parks (n = 76) and Biosphere Reserves (n = 40) within the Mediterranean bioclimatic region Country

National Parks (no.)

Albania Algeria Croatia Cyprus Egypt France Greece Israel Italy Lebanon Libya Morocco Portugal Spain Tunisia Turkey

3 6 5 1 0 2 10 3 15 1 3 7 0 5 6 9

Surfaces of National Parks (ha) 6,260 75,867 51,964 9,337 0 71,975 68,732 12,504 986,504 810 150,000 293,000 0 187,856 40,026 202,953

Biosphere Reserves (no.) 0 5 1 0 1 4 2 1 6 0 0 2 1 13 4 0

Area of Biosphere Reserves (ha) 0 152,806 200,000 0 7,000 296,585 8,850 26,600 256,135 0 0 975,4151 358 1,334,334 73,562 0

Source: Completed after Ramade (1997), Quézel and Médail (2003) and <www. unesco.org/mab/>, accessed 17 November 2008

TABLE 23.6. List of Ramsar sites within the Mediterranean bioclimatic region Country

Ramsar sites (no.)

Albania Algeria Croatia Cyprus Egypt France Greece Israel Italy Lebanon Libya Malta Monaco Morocco Portugal Slovenia Spain Syria Tunisia Turkey

3 39 1 1 2 5 10 2 42 4 2 2 1 24 17 2 36 1 1 11

83,062 2,897,115 11,500 1,585 105,700 136,552 163,501 366 55,489 1,075 83 16 10 272,010 73,784 955 156,351 10,000 12,600 157,782

206

4, 139, 536

TOTAL

Total area of Ramsar sites (ha)

Source: Modified from <www.ramsar.org>, accessed 1 April 2007 (see Chapter 9).

ranean but hardly more than 3 per cent is effectively protected, at least on paper. In fact most protected forested areas are too small and isolated to provide real security for habitats and their associated plant and animal communities. In some regions, for example in Greece, Croatia, Bosnia-Herzegovina, and Slovenia, almost 50 per cent of the land is still forested, sometimes with beautiful, nearly pristine forests as in parts of the Peloponnese and the Rhodope mountains in Greece. The Vikos-Aoos National Park in Epirus, northwest Greece (Figure 23.18), is still in a beautiful state of wilderness with Lynx, Bear, Wolf, Otter, Chamois, Roe deer, Wild boar, and Wild cat (Catsadorakis 2003). In contrast, no more than 1 per cent of French forests can be considered in a satisfactory state of conservation (Vallauri 2003). The situation in North Africa is even worse with an ongoing degradation and reduction of the forest cover without any effective conservation policy (Chapter 20). In contrast to vertebrates, which have enjoyed a long tradition of conservation, the protection and restoration of threatened plant populations is less advanced in the Mediterranean region, and many reintroduction programmes have unfortunately failed. This is probably due to a lack of preliminary integrative biological

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Fig. 23.18. The Vikos Gorge with its spectacular limestone cliffs and, in the middle distance, forested scree slopes and valley floor in the Vikos-Aoos National Park in Epirus, north-west Greece (photo: Jamie Woodward).

and ecological studies both on the target plants and their potential habitats. One interesting initiative for providing protection to threatened flora and endemic plant species (e.g. Antirrhinum pertegasii, Silene diclinis, Verbascum fontqueri) is the Plant Micro-Reserves (PMR) network of the Valencian region launched in 1994 by the Generalitat Valenciana (Laguna et al. 2004). This network of small (2–20 ha) reserves currently includes 207 PMRs which confer protection to 57 per cent of the critically endangered (CR sensu IUCN) plants present in this region. In France, several reintroduction programmes are carried out by official structures such as Conservatoires Botaniques Nationaux with some successful outcomes, notably in Corsica. Significant results have also been obtained within the framework of the European Union LIFE programmes, for example concerning the practical tools to restore populations of

a very rare west Mediterranean amphibious quillwort (Isoetes setacea) in temporary flooded pools threatened by the colonization of several shrub and tree species (Rhazi et al. 2004). A crucial point to ensure a rational conservation network is to identify biogeographic areas with both representative and high biodiversity value. Following the Important Bird Areas (IBAs) model developed by BirdLife International and covering 3,619 sites in fifty-one European countries, Plantlife International and IUCN (2003) have recently launched the Important Plant Areas (IPAs) programme in the Mediterranean region. The aim of this initiative is to identify areas that are appropriate for a site-based approach for conservation and to protect a network of the sites that are the most relevant for wild plants throughout the Mediterranean region. This network of IPAs should provide the framework for governments to achieve the Target 5

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of the Convention on Biological Diversity’s Global Strategy for Plant Conservation, which calls for the protection of 50 per cent of the most important areas for plant diversity. Three main criteria have been used to identify IPAs: (1) the site holds significant populations of species of global or regional concern (presence of threatened species); (2) the site has an exceptionally rich flora in a regional context in relation to its biogeographic zone (species richness); (3) the site is an outstanding example of a habitat type of global or regional concern (presence of threatened habitats). Some Mediterranean IPA projects are already completed, for example in Turkey (2002), and in Croatia (2003) where many of the eighty-eight identified IPAs lie outside the existing protected areas. Vertebrates, especially birds, provide many examples of a steady biodiversity recovery after restoration and proper management of habitats. This is the case in some wetlands in spite of an ongoing loss of marshland in the Mediterranean as a whole. In the Camargue, for example, the diversity and wealth of bird communities has dramatically increased in recent times as a result of an increase in the size of the populations of several species (herons, gulls, terns) and the appearance of new breeding species (e.g. the Spoonbill Platalea leucorodia, the Great white egret Egretta alba, the Glossy ibis Plegadis falcinellus, the Greylag goose Anser anser, the Cormorant Phalacrocorax carbo, and several others). Another spectacular example of biodiversity recovery is that of the Moulouya River mouth in northern Morocco (Brosset 1990). The size of this marshland has increased tenfold thanks to the building of large reservoirs in the upper course of the river and the expansion of irrigated farmland that raised the level of the water table. This water management strategy was initially introduced to improve local agriculture. It also indirectly increased the area of marshland and benefited many species so that the area is today one of the best breeding sites in North Africa for many rare species such as the Purple gallinule Porphyrio porphyrio, the Marbled teal Marmonetta angustirostris, and many others. Encouragingly, this example is a clear demonstration that some aspects of biodiversity can recover quickly as soon as favourable conditions reappear. Hundreds of species could be reinforced or locally reintroduced and many projects have been designed to preserve, reintroduce, or restore populations of particularly endangered species, including plant and animal breeds that had been selected by humans. For example, a project to reintroduce the Griffon vulture Gyps fulvus in the Cévennes, France, succeeded in the reestablishment—one century after its local extinction—

of a self-sustaining population of this species that today includes more than 200 individuals (Terrasse 1996). Other reintroduction projects for this species and for the rarer Black vulture Aegypius monachus elsewhere in southern France promise to be successful. The World Wildlife Fund launched several projects to reintroduce locally extinct populations of several species of vultures, especially the Black vulture Aegypius monachus in Greece (in the Dadia forest) and the Lammergeier Gypaetus barbatus in several countries of Mediterranean Europe. Among the mammals, the Sardous deer, a local population of the Red deer Cervus elaphus, has been successfully saved in Sardinia and recently reintroduced in Corsica. Several ongoing programmes contribute to reinforce endangered species such as the Iberian lynx Lynx pardus in Spain, as well as several species of mammals in Israel, e.g. the Mountain gazelle Gazella gazella. More stringent regulation policies and active action plans launched by BirdLife International for the protection of raptors succeeded in the restoration of many populations of endangered species (Muntaner and Mayol 1996). Reintroduction policies are faced with many problems, however. If properly executed on scientific grounds, they are extremely expensive and require clear strategies to raise public awareness. In addition, economic or political criteria often prevail over scientific arguments for biological conservation because it is much easier to raise funds for protecting flagship species to which public attention can be drawn than for enhancing the diversity of inconspicuous soil micro-invertebrates. In the semi-arid and arid south and east margins of the Mediterranean, successful projects currently include the restoration of the Oryx Oryx dammah and several species of gazelle and deer in several countries (e.g. Tunisia, Morocco, Jordan, Israel, and Syria). There is also currently much effort to preserve the myriad of ecotypes and gene pools of cultivated plants and domesticated animals that have been selected over millennia in all parts of the basin (Diamond 2002). One challenge is to identify and save these wild ancestral gene pools before they are lost for ever. International, governmental, and private organizations currently combine their efforts to preserve these genetic resources of old breeds of cattle, sheep, goats, and the Mediterranean buffalo. In the same way, many programmes aim to conserve genetic resources of fruit trees, grapes, field crops, forage species, vegetables, and ornamental plants (Charrier 1995). In this context, one promising initiative is the International Centre for Advanced Mediterranean Agronomic Studies (CIHEAM) which includes several institutes in Italy (Bari), Greece (Chania), France (Montpellier), and Morocco (Meknès). These Institutes

Biodiversity and Conservation

carry out research involving both in situ and ex situ study of local races of traditional Mediterranean plants. One aim of these institutes is to help promote sustainable development programmes using these local ecotypes.

The Uncertain Consequences of Global Change for Mediterranean Biodiversity Most of the ups and downs in Mediterranean living systems are closely linked with human population pressures which have changed many times through the long common history of ecological systems and human societies. The dynamics of human populations will be a major factor in the future of the Mediterranean. The northern countries which contributed about two-thirds of the human population in the Mediterranean basin in the 1950s will contribute only about one-third in 2025 as a result of demographic growth in the southern and eastern countries of the basin. By 2025 in these countries, from Morocco to Turkey, the human population will have become five times that of 1950 (Grenon and Batisse 1989). In coastal areas which are particularly sensitive to disturbance, the projected changes in human population are 200 million in 2025 instead of 145 million in 2000, with 170 million instead of 95 million living in urban areas. In this context, the future of biological diversity will depend on how humans societies will learn to live together and with their natural heritage in the forthcoming decades. The Mediterranean is the most popular tourist destination in the world, which means that conservation and proper management of coastal areas in particular will be one of the most serious challenges in the forthcoming decades (Chapter 13). This will be especially important on the 18,000 km of coast on the Mediterranean islands, all the more so because in the future, tourism will be the major resource in many countries (Benoit and Comeau 2005). Therefore, in addition to natural resource conservation and management, priority should be given to the protection of coastal areas and wetlands where the strongest pressures occur. The various components of global change, especially climate changes, are other challenges to meet. Scenarios of the IPCC (2001) show that the Mediterranean region will be particularly affected by global warming and a decrease in rainfall. Water shortage in the future could be critical in a region with increasing demand for water supply for an ever-growing population in large cities and the enormous impact of tourism

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(Chapter 21). In addition, deforestation and increasing rates of runoff due to deforestation and soil erosion in North Africa will accelerate the process of desertification (Chapter 20). With an expected temperature increase of c.3–5◦ C in the Mediterranean during the twentyfirst century, potential evapotranspiration should on average reach 200 mm annually, which is equivalent to a loss of 50 mm in annual rainfall (Le Houérou 1990). The expected shifts in vegetation belts resulting from increased aridity will be an upward move of 80–100 m in altitude and 50–80 km northwards in latitude. As a whole, the geographical distribution of the natural vegetation will undergo a shift towards more thermophilous types. On the other hand, if the predictions are accurate, productivity will increase as a result of the extension of the growing season by one to two weeks (Penuelas et al. 2002), and the reduction in winter cold stress. Furthermore, with a doubling of livestock between 1950 and 1986 and a current density of about one sheep-equivalent per hectare against an overall carrying capacity of one sheep-equivalent per 10–12 ha, there will inevitably be a severe encroachment of desert-like conditions over the entire arid zone of the southern side and most of the eastern side of the Mediterranean. In addition, the expected rise of the sea level by c.40 cm during the twenty-first century (IPCC 2001) will threaten many coastal areas and the millions of people who live there (Chapter 13). A hierarchy of the expected impacts of the drivers linked to global change upon the ten Mediterranean regional hotspots of biodiversity (Médail and Quézel 1997, 1999), was assessed by the group of RICAMARE experts (Table 23.7) (Troumbis et al. 2001). Southern hotspots will be threatened by the extension of arid conditions, whereas northern hotspots will be endangered by an increasing impact of some disturbance events (fires), albeit favoured by the northward extension of Mediterranean-type ecosystems. The sensitivity of habitats will differ according to species assemblages, and a complex pattern of change will occur, depending on location, species composition, and the magnitude of trophic interactions. Saxicolous communities will probably be less affected than grassland, forest, and coastal communities. Fortunately, many rare plant endemics are located in communities with high stress levels (e.g. cliffs, rocks, and screes), which constitute the best refugia against climate changes (Figure 23.19) (see Médail and Diadema 2009, for a recent review). On the other hand, rocky grasslands and steppes could be seriously endangered by the direct ecological consequences of global change including land use changes, aridification processes, increases in the productivity and biomass of

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TABLE 23.7. Major influences on biodiversity in the Mediterranean hotspots shown in Figure 23.1 according to the group of RICAMARE experts. Where data are available, the relative impact of each factor is given in each case. Note that the Betic-Rifan complex has been divided into two groups in this analysis because of marked contrasts in land use practices. Mediterranean hotspots High and Middle Atlas Rif Mountains Betic region (Andalucia) Maritime and Ligurian Alps Western Mediterranean Islands Southern and Central Greece Crete Southern Anatolia and Cyprus Syria, Lebanon, and Israel Mediterranean Cyrenaica Djurdjura-Kabylies

FIRE

URBA

FORE

GRAZ

AGRI

INVA

N DEP

CO2

UV B

TEMP

RAIN

2 2 3 2 3 3 2 2 3 2 2

0 1 2 2 2 2 1 2 3 2 1

1 2 1 1 1 2 1 1 2 3 2

3 1 1 1 2 2 2 3 1 3 2

1 3 2 1 1 2 2 2 1 2 2

1 1 1 1 2 2 2 1 0 2? 1

? 1 ? ? ? ? ? ? ? ? ?

? ? ? ? ? ? ? ? ? ? ?

2? 1? 1? 2? ? ? ? 2? ? 1? ?

2 2 3 2 2 3 3 1 3 2 3

3 2 3 2 2 3 3 1 3 2 2

Notes: Drivers: FIRE: impact of fires; URBA: urbanization; FORE: forestry practices; GRAZ: grazing; INVA: biological invasions; N DEP: nitrogen deposition; CO2 : CO2 deposition; UV B: ultra-violet deposition; TEMP: putative sensitivity to global warming; RAIN: putative sensitivity to modifications of the rainfall regime. Scale of driver impacts: 0 = no impact; 1 = low impact; 2 = medium impact; 3 = strong impact. Source: Modified from Troumbis et al. 2001. RICAMARE is ‘Research in Global Change in the Mediterranean: a Regional Network.

shrubs and trees, and areal extensions of severe competitive aliens. The question of the extent to which organisms will be able to cope with the new challenges generated by climate change is still largely open, especially

because climate change appears to be taking place at an unprecedented rate in recent human history. In the Mediterranean region there is recent evidence that microevolutionary changes can occur rapidly in fitnessrelated traits such as the breeding season of birds (Visser

Fig. 23.19. Cliffs are important habitats for several rare birds as well as rare endemic plant species throughout the Mediterranean basin (photo: Jacques Blondel).

Biodiversity and Conservation

et al. 2003) or the flowering time in plants (Penuelas et al. 2002). But differential responses of organisms engaged in complex food chains or symbiotic associations may disrupt interactions that are essential for ecosystem functioning such as pollination or seed dispersal. Indeed, different types of responses to global change must be incorporated into management decisions related to conservation and restoration ecology in order to allow endangered populations to adapt themselves to these changing selective pressures (Stockwell et al. 2003). Studying the mechanisms whereby organisms will become adapted to their environment is an issue of major importance for predicting their response to ongoing climate changes. The combination of genetic

and palaeoecological studies has recently provided several interesting insights into the capacity of plants to cope with drastic climatic changes. These studies have also demonstrated the crucial role of glacial refugia for the long-term persistence of the Mediterranean biodiversity (e.g. Tzedakis et al. 2002; Petit et al. 2003). A recent detailed analysis based upon intraspecific phylogeographical studies of eighty-two plant species led to the identification of fifty-two refugia located in the Mediterranean region (Médail and Diadema 2009). These ecologically and climatically stable areas have a high conservation priority because they represent significant reservoirs of endemic plants and of unique genetic diversity for Mediterranean species. But these refugia

Established refugia

(a) Refugia 100 km

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100 km

Boundary of Mediterranean region

(b) Population density N

Inhabitants per km 2

> 1,000 250 –1,000 100 – 250 50 – 100 0 – 50

Fig. 23.20. Distribution maps within the Mediterranean bioclimatic region (limits indicated by a broken line) of the fifty glacial refugia based upon (a) the phylogeographic structure of eighty-two plant species, and (b) of the human population density, following a 100 × 100 km lattice (after Médail and Diadema 2006).

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constitute some of the most threatened territories of the Mediterranean basin, owing to higher than average demographic pressures: c.25 per cent of the areas where refugia occur are located in highly populated regions (>250 people per km2 ) (Médail and Diadema 2006, 2009; Figure 23.20). After 10,000 years or more of cohabitation between humans and nature, most Mediterranean ecosystems are so inextricably linked to human interventions that the future of biological diversity cannot be disconnected from that of human affairs, which have too often been characterized by bitter conflicts. Whatever the future of people and their environment, landscape conditions will continue to be shaped and driven by the long common history they share (Foster et al. 2003). Thus, the greatest challenge for the decades to come is to promote sustainable development through the control of population growth and the development of appropriate indicators to assess how environment and development should be integrated. The Mediterranean region, at the crossroads of three continents, between developed countries in the north and developing countries in the south and east, between deserts and fertile lands, has always been and still is an area of both division and convergence. Because it is, in a sense, a microcosm that is representative of many worldwide problems, it is an exceptional laboratory and pilot region for launching a strategy for sustainable development on a regional scale.

Acknowledgements We are deeply indebted to Jamie Woodward who invited us to write this chapter. We also thank Oxford University Press for giving us this opportunity to make a point on the state of biodiversity in this fascinating region. Jamie Woodward and Kathy Willis much improved a first draft of the manuscript. Thanks to them.

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Index Bold numbers denote reference to illustrations or tables and their respective captions abstraction 79 groundwater 571, 590, 591 see also aquifers lakes 260, 274, 276, 278, 279, 281 rivers 241, 345, 347, 523 Adriatic Sea 11, 16, 33, 35, 37, 39, 42, 44, 51, 52, 53, 56, 57, 75, 77, 78, 84, 103, 197, 290, 301, 368–9, 385, 387, 388, 392, 401, 406, 472, 496, 497, 508, 515, 516, 522, 531–2, 600, 622, 635, 636 Aegean arc 10, 13, 456, 473, 474 Aegean Sea 8–13, 15, 18, 21–2, 33, 34, 35, 37–9, 42, 44, 51, 52–3, 56, 57, 74, 75, 78, 103, 123, 154, 245, 246, 269, 338, 347, 385, 387, 388, 390, 398, 399, 400, 401, 403, 406, 421, 435, 436, 450–1, 454, 456, 459, 470–1, 472, 478–80, 482–3, 485, 493, 496, 499, 500, 503, 504, 508, 509–510, 515, 534, 566, 579, 603 Aegean volcanic province 436, 450–57 see also volcanoes and volcanism aerosols 415–7, 418–19, 437, 439–40, 449–50, 529, 599–601, 607–11 see also aeolian processes and landforms; TOMS aerosol haze see air quality Aeolian Islands 436, 446–9 see also volcanoes and volcanism aeolian processes and landforms 415–25 dunes and other landforms 325, 386, 405, 422–5, 571 dust sources and trajectories 58, 75, 119, 175, 415–22, 599 see also TOMS; loess 420–2 loess 176, 304, 415, 416, 420–2, 425 see also aeolian processes and landforms afforestation 208, 249, 571 see also reafforestation African plate 8, 11, 36, 324, 435, 450, 469 see also tectonics and landscape development aftershocks see and seismicity earthquakes aggradation (fluvial) 17–18, 322, 323, 327, 329, 331–5, 338, 339 see also river systems and environmental change

agriculture see also pastoralism; irrigation; land degradation fertility and nutrients 573, impacts 182, 184, 192, 239, 272, 571, 586, 603, 609, 617, 621 see also soils land use change 224–5, 250, 263, 272, 276, 301, 314, 406, 423, 572, 575, 621, 625, 633 mechanization and intensification 179, 564, 618, 637–8 origins and early expansion 213–4, 222–3, 271, 338, 546, 617, 638 ploughing 182, 222, 239, 249, 336 see also tillage pollution 274, 280, 591 see also eutrophication storm damage 534, 535 volcanic hazards and agriculture 437, 438–9, 455 agricultural terraces 214, 239, 301, 303, 337, 591 abandonment 214, 224, 549 see also land abandonment Aguas River, Spain 24, 25, 26, 245, 328, 330, 333, 334 air masses 37, 39, 42, 44, 47, 48, 55, 73, 513–4, 519, 537, 601–3, 605 see also climate air quality 437, 440, 553, 561, 599–600, 603, 605, 609, 610–611 aerosol haze 437, 449, 601, 608–10 see also aeolian processes and landforms climatic controls on 600–3 climate and the water cycle 610–11 oxidation capacity 607–8 ozone dynamics 604–6 upper tropospheric pollution plume 603–4 Albania 11, 75, 233, 236, 256, 257, 263, 266, 275, 277, 355, 357, 365, 367, 378, 386, 448, 495, 497, 508, 541, 587, 588, 594, 596, 631, 640 Alboran basin 9, 10, 385, 387, 389 Alboran Sea 10, 13, 39, 40, 41, 42–3, 58, 103, 105, 120, 121, 122, 377, 386, 388, 495–6, 508, 510, 532 Alcantara River, Spain 335

Aleppo pine 206, 207, 210, 220, 237, 551 Alexander the Great 485 Alexandria 40, 69, 70, 71, 386, 389, 418, 424, 471, 498, 499 Alfios River, Greece 239, 320, 321, 346 Algarve 211 Algeria 10, 19, 20, 39, 40, 41, 42, 58, 78, 108, 116, 150, 197, 232, 233, 236, 244, 247, 274, 275, 287, 315, 319, 376, 386, 416, 417, 418, 424, 469, 470, 480, 481, 482, 485, 495, 496, 513, 514, 515, 516, 517, 532, 533, 535, 541, 571, 587, 589, 590, 591, 592, 594, 595, 596, 618, 626, 628, 631, 640 Algerian basin 7, 9, 10, 41, 385, 387, 388 Algerian current 40, 41 Algiers 41, 71, 74, 495, 536 Alkyonides Gulf 22–3 allopatric speciation 140, 148, 160, 216 alluvial archaeology 319 see also geoarchaeology alluvial fans 7, 18, 23, 177, 178, 325, 359 alluvial history see river systems and environmental change alluvial soils 234, 332, 177, 178 Almanzora River, Spain 249 Almeria, Spain 7, 19, 40, 41, 70, 109, 187, 245, 567 Alpine-Himalayan belt 5, 6 Alps (Austrian/French/Swiss) 10, 11, 12, 38, 70, 74, 77, 84, 113, 115, 124, 141, 143, 149, 169, 174, 185, 190, 191, 195, 255, 263, 293, 301, 325, 331, 344, 361, 370, 373, 376, 377, 417, 418, 421, 515, 521, 531, 585, 600, 603 Alps-Betics orogen 7, 10 AMS radiocarbon dating see radiocarbon dating Anatolia 8, 11, 75, 86, 116, 139, 141, 144, 156, 169, 171, 177, 207, 208, 215, 222, 255, 261, 262, 263, 385, 390, 403, 469, 515, 583, 616, 621, 644 Anatolian Fault, North (NAF) 470, 472–3, 474, 477, 478, 480

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Index

Anatolian plate 473 Ancient Greece 469 see also Classical Period Andalucia 214, 264, 265, 516, 521, 567, 572, 630, 644 Andorra 360, 516, 522 anoxia 269 see also sapropels Antarctica 73 Antarctic gateway 91 Antarctic ice cores 124 Antarctic ice sheets 91, 92, 95, 126 anticyclones 47, 74, 77, 78, 126, 127, 418, 519, 521, 531, 537, 602, 603, 604, 605, 610 see also cyclones and cyclogenesis; air masses; climate aphelion 45, 46 Appenines, Italy 16, 357, 369–370 Apulia, Italy 10, 139, 288, 295, 385, 400, 401, 508 aquifers 196, 251, 274, 276, 281, 287, 295–8, 299, 307, 309, 313, 571, 583, 584, 586, 590, 591, 592 see also water resources Arabian Desert 415, 418 Arabian Peninsula 602, 603 Arabian plate 36, 144, 147, 324 see also tectonics and landscape development Arabian Sea 104 archaeology 140, 281, 319, 337 see also geoarchaeology architecture 306, 313, 315, 485 Argive plain, Greece 338 aridification 93, 94, 95, 643 see also aridity; aeolian processes and landforms aridity 47, 57, 93, 97, 99, 100, 101, 102, 107, 114, 116, 119, 123, 126, 127, 169, 171, 178, 204, 209, 211, 234, 295, 376, 422, 568, 583, 585, 621, 643 Arles, France 344, Arno River, Italy 345, 388, 404, 621 aromatic plants 204, 206 artefacts 476 see also coins; pottery; archaeology; stone tools Asia 69, 73, 75, 93, 107, 112, 113, 139, 141, 142, 144, 145, 146, 147, 155, 209, 216, 229, 425, 483, 510, 514, 593, 602, 603, 604, 606, 610, 629 see also Eurasia Asia Minor 176, 469, 522, 593 Asia Minor Current 41 Asian monsoon 39, 46, 55, 75, 604 see also Indian monsoon astronomical parameters see Milankovitch cycles Aswan High Dam 39, 40, 97, 232, 272, 405, 585, 590 see also Nile River Athens 69, 70, 71, 470, 476, 479, 484, 503, 514, 523, 535

Atlantic Ocean see North Atlantic Atlas Mountains 21, 58, 70, 75, 108, 116, 123, 149, 150, 154, 177, 207, 212, 215, 255, 261, 269, 272, 273, 287, 303, 354, 360, 361, 375–376, 416, 517, 616, 628, 644 atmospheric instability see cyclones and cyclogenesis; storms and floods; climate Australia 69, 157, 305, 541, 546, 547, 579, 615, 629 Austria 11, 74, 301 Axios River, Greece 245, 246, 347, 388, 621 Azores 78, 440 Azores High 37, 386, 601–602 back-arc basins 7, 8, 9, 10, 11, 13, 331 see also tectonics and landscape development badlands 5, 7, 18–20, 24, 25, 26, 185, 187, 239, 320, 334, 570 Bagni de Tivoli, Italy 306 Balearic basin 20, 43, 385, 387, 388, 400 Balearic Islands 10, 423, 424, 496, 513, 514, 515, 516, 532, 533, 536, 623, 631 see also Majorca Balkan Peninsula 102, 111, 139, 141, 142, 233, 287, 301, 513, 515, 525, 584, 600, 638 Balkans 74, 75, 78, 80, 111, 143, 144, 149, 156, 159, 176, 223, 233, 255, 261, 264, 266–7, 274, 331, 355–7, 362, 365–9, 470, 522, 596, 635 Barcelona 86, 282, 360, 516, 519, 522, 532, 533, 534, 572 basalt 14, 16, 171, 335 baseflow 586 see also runoff base level 7, 15, 16, 18, 19, 23, 24, 25, 26, 287, 293, 296, 301, 304, 307, 325, 334–5, 389, 480, 483 see also tectonics and landscape development Basento River, Italy 388 basin and range environments 7, 17 see also tectonics and landscape development beaches see coastal environments bed load see sediment loads and yields bedrock erosion and weathering 17, 171, 173, 175, 185, 189, 235, 290, 292, 298, 302, 303, 306, 311, 325, 327, 357, 378, 395, 484 Benioff zone 8, 13 Betic Cordillera, Spain 5, 7, 10, 11, 24, 293, 333, 616 Betic corridor 7, 14, 36, 94, 389 see also Rifian corridor bioclimate and bioclimatic zones 127, 149, 151, 203, 205, 630, 640, 645

bioconstruction 394, 396, 397 biodiversity 140, 150, 157–60, 203, 206, 214–7, 224, 225, 250, 258, 266, 272, 274, 277, 279, 281, 282, 396, 406, 596, 615–646 see also species diversity biodiversity and conservation 615–646 see also conservation biodiversity decline 629–40 conservation frameworks 640–3 fire and biodiversity 624–6 global change and biodiversity 643–6 habitat destruction and fragmentation 618–24 hotspots 615–6, 644 human impact on 616–29 invasive species 626–9 biogeography 5, 142, 149, 155, 210 biomass 93, 123, 158, 173, 209, 216, 396, 546, 556, 573, 574, 575, 576, 579, 603, 609, 643 biostratigraphy 401, 498 see also bioconstruction; pollen and pollen records; foraminifera; diatoms bioturbation 36, 37, 182, 267 bird fauna 125, 140, 141, 142, 143, 147–8, 151, 152, 154, 155, 156, 214, 266, 274, 276, 281, 618, 621, 623, 624, 625, 628, 633, 634, 636, 637, 638, 641, 642, 644 BirdLife International 641, 642 Black Sea 10, 15, 23, 39, 44, 74, 75, 108, 247, 255, 274, 275, 355, 402, 522, 604, 609 black shales 33, 36 blue-green algae 267 Blue Plan 232 Bora wind 35, 37, 44, 74, 77, 78, 513, 531, 537 see also climate; storms and floods boreal summer 45, 46, 54, 99, 100, 101, 102, 116, 117, 123, 127 boreal winter 45, 46, 54 Bosnia-Herzegovina 11, 275, 288, 353, 355, 541, 587, 588, 594, 640 boulder bed channels 309, 324, 343, 344 see also gravel bed rivers braided rivers 18, 345, 404, 422 see also gravel bed rivers breccia 456 Bronze Age 18, 147, 158, 185, 192, 251, 269, 270, 271, 315, 336, 338, 339, 398, 451, 452, 460, 499, 500 browsing 212, 220, 221, 272 see also grazing Brown bear (Ursus arctos) 142, 143, 144, 146, 634 Bulgaria 11, 108, 111, 143, 275, 443, 444, 550, 609 Butzer, Karl 336, 337

Index Büyük Menderes River, Turkey 18, 263, 388 Byzantine Period 272, 503 see also Early Byzantine Tectonic Paroxysm C3 93, 98, 127 C4 93, 98, 127 Cairo 74, 389, 416, 418 Calabria 7, 10, 11, 12, 13, 97, 113, 119, 141, 171, 179, 190, 390, 398, 448, 484, 495, 496, 497, 508, 514, 535 Calabrian arc 5, 7, 9, 12, 13, 16, 385, 386, 387, 390, 391, 398, 500 see also subduction; tectonics and landscape development Calabrian pine 206, 207, 210, 220 calcium carbonate 53, 93, 98, 100, 115, 124, 127, 128, 171, 178, 185, 211, 258, 263, 264, 266, 267, 268, 287, 288, 290, 292, 293, 296, 297, 298, 299, 301, 305–6, 308, 312, 313, 315, 333, 393, 394, 423, 424, 442, 445, 476, 586 see also calcrete; limestones; travertine; tufa; speleothems calcrete 24, 127, 178, 244, 305, 333 see also cemented till; cementation Calderone glacier 80, 357, 362 California, USA 69, 157, 223, 440, 541, 546, 547, 553, 554, 615, 627, 628 Camargue wetlands, France 213, 263, 274, 621, 623, 629, 642 Campanian volcanic province 436, 442–6 see also volcanoes and volcanism Canary Islands 148, 154, 155, 207, 630 Cantabrian Mountains, Spain 108, 123, 143, 353, 354, 359 see also Picos de Europa carbon and the carbon cycle 91, 93, 127, 197, 416, 599, 608 see also carbon dioxide inorganic carbon 197, 599, 603, 608, 609, 611 organic carbon 53, 93, 127, 197, 567, 574 carbon dioxide 80, 91, 93, 107, 112, 197, 217, 222, 223–4, 290, 297, 303, 437, 439, 440, 442, 449, 599, 603, 644 Carboniferous limestones 295 Carpathian Mountains 7, 11, 13, 141 Catalonia 19, 85, 97, 109, 190, 217, 248, 423, 514, 516, 517, 521, 522, 525, 526, 528, 530, 531, 532, 533, 535, 536, 541, 546, 547, 549, 552, 556, 596 catastrophic flooding 74, 122, 194, 323, 442, 515, 518, 519, 520, 522, 523,

525, 526, 530, 531, 535, 537 see also storms and floods Caucasus Mountains 10, 108, 149, 255, 298, 378, 435, 436 Caves and cave deposits 56, 102, 103, 104, 123, 141, 146, 154, 156, 157, 174, 249, 287, 288, 290, 293, 295, 296, 297, 298–300, 301, 305, 307, 309, 310–13, 314, 315, 357, 390, 393, 394, 395, 402, 403, 415, 421, 422, 443, 444, 451, 484, 571, 633, 635–6 see also rockshelters; speleothems; karst Cave salamander (Proteus anguinus) 157–8, 636 Cave bear (Ursus spelaeus) 146 cemented till 365, 368, 370 see also cementation; calcrete cementation, 179, 292, 298, 303, 305, 306, 308, 321, 322, 325, 332, 333, 334, 393 see also calcrete; cemented till Cenozoic Era 79, 91, 92, 93, 94, 98, 104, 255, 295, 307, 324, 445 see also global cooling cereals 240, 618, 633 Chad, Lake 55, 418 chamois (Rupicapra rupicapra) 640 channel incision 7, 15–6, 18, 22–6, 195, 295, 301–2, 306, 319, 321–2, 327–8, 330, 333–5, 338–9, 345–6, 367, 389, 481, 483, 529 see also river systems and environmental change; gully development channel management see river systems and environmental change charcoal 102, 112, 206, 208, 223, 270, 312, 332, 544, 546, 553, 617 check dams 239–244 see also dams; reservoirs chemical weathering see weathering and rock breakdown; dissolution Chile 69, 157, 541, 615 chotts 261, 262, 405, 416–7, 421, 424 chronosequence 174 circumpolar vortex 119 see also jet stream cities see urban environments and urbanization Classical Period 272, 306, 319, 335, 403, 435, 448, 449, 503, 564, 565 see also Roman Period; Bronze Age; Ancient Greece clays and clay minerals 19, 20, 51, 54, 172, 173, 174, 178, 185, 189, 272, 297, 402, 416, 420, 422, 425, 439 clearance, woodland 158, 192, 208, 271, 272, 336, 563, 572, 574, 575, 577, 618 see also deforestation

653

climate: air temperature 69–71, 74, 77–81, 83–4 cyclones and cyclogenesis see storms and floods future climate 80–5 heatwaves 73, 77–9 hurricanes (‘Medicanes’) and tornadoes 532–5 modelling 37, 47, 54, 55–6, 58,83–6, 91, 94, 99, 103, 169, 176, 223, 232–3, 515, 519–23, 542, 554, 555, 576–7, 603, 605–6, 610–1 orographic influence on 35–6, 57, 93, 103, 515, 521–2, 532, 534, 603, 605 rainfall 19, 39, 54, 69–75, 79–80, 82–86, 99–102, 120–3, 126–127, 171, 179–180, 183–191, 204, 212, 220, 229–245, 265, 290, 295, 303, 309, 344, 361, 425, 439, 445, 513–37, 548, 552, 555, 568–71, 574, 576, 583–5, 592, 602, 610–11, 643 seasonality 37, 39, 45–7, 55, 71, 73–5, 80, 85, 89, 91, 93, 97, 99–102, 117, 126–7, 178, 204, 229, 232–6, 247, 249–50, 264, 267, 276, 295, 417, 523, 532–3, 546, 548 snowfall 69–70, 91, 150, 174–5, 190, 268, 296, 301, 303, 354–78, 439, 527, 534–5 see also snowmelt; snowline winds 38–9, 40, 42–44, 55, 69–70, 74–5, 77–8, 93, 119, 179, 386, 388, 415–25, 457, 513–5, 521, 531–35 see also storms and floods; Bora; Etesian; Levante; Mistral; Tramontana climate change see climate; Quaternary environmental change; Heinrich Events; Dansgaard-Oeschger events; Milankovitch cycles; Little Ice Age; Medieval Warm Period coastal environments 385–407 see also sea-level change aquifers 307, 309, 346 beaches 386, 393, 398, 402, 405–6 biology 395–7 caves 102, 312, 394, 395, 403 coastal flooding 404, 406, 527 currents 40–4, 386, 388, 389 deltas 15–16, 18, 40, 83, 95, 151, 214, 263, 272, 274, 277, 319, 336, 346, 347, 386, 388–9, 393, 402–6 dunes 386, 405, 422, 423–4 harbours 77, 394, 395, 398–400, 402–4, 405, 471, 477, 480, 482, 493, 498, 501

654

Index

coastal environments (cont.) Holocene development 401–2 see also deltas littoral cells 386–88, 389 plains 77, 97, 109, 197, 319, 325–7, 338, 344, 390, rocky shores 23, 222, 287, 288, 289, 315, 386, 395, 396 tectonics 23, 36, 385–93, 398–400 see also tsunamis wetlands 83, 95, 97, 213, 255, 257–60, 263, 272–3, 275, 278, 279, 281 winds 77, 78 coins 321, 476 see also artefacts colluvial processes see hillslopes Como Lake, Italy 255, 258, 262, 273, 278 coniferous forest 95–7, 117–8, 127, 141, 203–5, 207–8 see also forest communities conservation 126, 139, 145, 160, 206, 212, 213, 225, 237, 239, 245, 255, 260, 267, 272–4, 276, 279–80, 281, 312–4, 563, 564, 565, 571, 572, 578–9, 583, 592, 615, 621, 624, 627, 632, 637, 640–3, 645 see also biodiversity and conservation Constantinople 504 see also Istanbul Corinth, Gulf of, Greece 7, 21, 22, 23, 390, 400, 401, 471, 479, 480–1, 484, 485, 495, 496, 502–3, 505, 506, 507, 508, 510 Corinth, Greece 20, 470, 486 Corsica 10, 11, 40, 158, 206, 232, 323, 324, 329, 344–5, 353, 354, 370, 373, 416, 417, 420, 421, 424, 516, 572, 624, 626, 628, 631, 632–3, 635, 641, 642 Cosquer cave, France, 141, 394, 395, 633 Côte d’Azur, France 358, 400, 496–7, 505, 508, 516–7, 535 crayfish 257, 272–3, 278–9, 628 Cretaceous 11, 16, 93, 127, 196, 293, 307 Crete 21, 35, 39, 113, 151, 154, 175, 206, 207–8, 212–5, 216, 220, 221, 224, 249, 263, 293, 323, 328, 330, 338, 340, 343, 344, 345, 354, 355, 397, 399, 400, 420, 421, 422, 424, 441, 450, 454, 460, 471, 479, 480–3, 485, 487, 493, 495, 496, 497–501, 505, 530, 547, 572, 575, 600, 603, 607, 608, 609, 616, 631, 632, 644 Croatia 11, 77, 143, 169, 261, 274, 275, 288, 293, 296, 298, 301, 306, 312, 314, 353, 369, 385, 386, 532, 535, 543, 587, 588, 594,

631, 636, 640, 642 see also Dalmatia cryosphere 91, 92, 319, 353–78 see also glacial and periglacial environments cryptic refugia, see refugia cultivation see agriculture cultivars 160, 638–9 see also agriculture; domestication cyclones and cyclogenesis see storms and floods; air masses Cyprus 10, 11, 33, 39, 41, 42, 70, 75, 154, 169, 207, 215, 233, 236, 260, 263, 275, 385, 386, 391, 400, 401, 418, 424, 485, 495, 496, 500, 504, 505, 508, 515, 534, 537, 587, 588, 594, 595, 596, 616, 621, 631, 634, 638, 640, 644 Cyrenaica 215, 255, 324, 325–327, 616, 630, 644 see also Libya; Tripolitania Dalmatia 78, 143, 288, 309, 422, 498, 531, 636 see also Croatia dams (natural and artificial) 20, 39, 40, 86, 97, 193, 194, 232, 239–44, 247, 249, 251, 257–60, 272, 274, 276, 279, 281, 314, 337, 345, 346, 367, 369, 405, 471, 481, 483, 485, 525, 527, 530, 583, 585, 589, 590, 591, 592 see also reservoirs; check dams Dansgaard-Oeschger events 45, 57–8, 331 see also sub-Milankovitch climate variability dating methods see lichenometry; luminescence; orbital tuning; potassium-argon; radiocarbon; uranium-series Dead Sea 56, 79, 102, 103–4, 105, 119, 255, 257, 261, 265, 266, 268, 272, 276, 321, 473, 474, 476, 505 see also Lisan, Lake debris flows 18, 20–1, 193, 437, 439, 445, 500, 505, 535 see also mass movements deciduous forest 92, 95–8, 100–101, 107, 109–10, 115–7, 127, 151, 173, 203–6, 216, 222, 271–2, 556, 572, 623–4, 626 deep water 34, 35, 37, 40, 42–4, 52–3, 94, 98, 99, 119, 122 see also Mediterranean Sea; North Atlantic Deep Water (NADW) deflation 175, 262, 305, 415, 423–5, 571 see also aeolian processes and landforms deforestation 80, 237, 239, 270, 272, 273, 274, 281, 314, 415, 425, 522, 536, 571–2, 575, 591, 615, 618, 632, 643 see also clearance

deglaciation 268, 375, 392 see also glacial and periglacial environments dehesas 158, 206, 225, 617 deltas see coastal environments desalination 70, 587 deserts see aridity; aeolian processes and landforms desertification 204, 220, 249, 415, 564, 565–6, 579, 615, 618, 643 see also aridification diatoms 123, 267, 268, 270, 273, 279 Dinaric Alps 39, 149, 255, 290, 314, 353, 354, 357, 368, 531–2, 583 Dinaric karst 157, 296, 301, 303, 304 dissolution 37, 171, 175, 179, 185–6, 258–9, 261, 264, 287, 290–2, 305, 314, 396, 416 see also weathering; solute loads DNA 125, 139, 142, 216 domestication 154, 157, 160, 220, 563, 638, 639, 642 drought 37, 69, 72, 73, 77, 79, 83–86, 100, 103, 113, 117, 123, 126, 128, 173, 204, 209, 218, 223, 249, 259, 265, 269, 319, 531, 535, 537, 548, 550, 555, 564, 568, 572, 592, 600, 610, 628, 631, 632 see also climate drylands see aridity dunes see aeolian processes and landforms; coastal environments dwarfism in mammals 154, 157, 634 Early Byzantine Tectonic Paroxysm 400, 471, 477 see also earthquakes and seismicity earthquakes and seismicity 5, 7, 11–3, 20–3, 190, 325, 390, 399–400, 401, 437, 444, 446, 450, 452, 457, 459, 469–87 see also earthquakes and tsunamis; tectonics and landscape development in Antiquity 471, 474–6, 481–4 data sources and methods 474–6 hazard management 485–7 impacts 478–84 earthquakes and tsunamis 493–510 see also coastal environments Eastern Mediterranean Deep Water (EMDW) see Mediterranean Sea Ebro delta, Spain 18, 388, 402, 404, 405 Ebro River basin, Spain 15, 16, 20, 39, 151, 186, 250, 262, 265, 329, 360, 389, 402, 416, 425, 516, 527, 530, 569, 585 eccentricity see Milankovitch cycles ecosystems 83, 90, 150, 151, 157, 159, 203–225, 234, 237, 249–50, 267, 269–70, 274–81, 287, 314, 378, 415, 439, 541, 543, 546, 550–53, 567–9, 575, 578, 586, 590, 599,

Index 615–8, 621, 624–8, 638, 643, 645, 646 Egypt 10, 57, 58, 74, 77, 78, 84, 145, 213, 258, 269, 272, 274, 275, 277, 287, 288, 293, 295, 301, 305, 307, 386, 389, 400, 401, 415, 418, 422, 424, 425, 471, 477, 480, 483, 485, 508, 583, 586, 587, 588, 590, 593, 594, 595, 596, 631, 632, 640 El Niño 73, 574 endemism 125, 140, 143, 146, 148, 149, 150, 151, 154–9, 203, 211–2, 215–6, 266–7, 274, 279, 282, 615, 622, 623, 626, 628–36, 641, 643–5 see also biodiversity Eocene 91–7, 98, 144, 291, 300 ephemeral streams and lakes 17, 232, 234, 244–5, 247–8, 250, 258, 261, 263, 277, 305, 325, 421, 422, 523, 525, 527, 585 see also river systems and environmental change Ephesus, Turkey 338 Epirus, Greece 19, 109, 151, 240, 321, 322, 367, 640, 641 Equilibrium Line Altitude (ELA) 354, 358, 362, 363–4, 366, 369–70, 373 see also snowline; glacial and periglacial environments erosion see soils; hillslopes; incision essential oils 204, 218–20 Etesian wind 74, 77, 78 see also climate; storms and floods Etna, Sicily 13, 14, 17, 172, 180, 335, 354, 357, 435, 436, 438, 441, 449–50, 458, 459, 460 see also volcanoes and volcanism Euboea, Greece 338, 397 European Commission (EC) 313, 544, 550, European Union (EU) 86, 224, 241, 250, 267, 280, 487, 535, 536, 564, 575, 599, 605, 609, 635, 636, 637, 641 Eurasia and Eurasian plate 5, 10–11, 14, 36, 45, 94, 139, 144–6, 147, 255, 293, 324, 435, 469, 472, 474, 563 eutrophication 257, 259, 260, 266, 274, 276, 277, 280–1, 313, 591, 599, 635 see also lakes and wetlands Euphrates River see Tigris-Euphrates evaporation 5, 14, 33, 37, 39–43, 45, 48, 52, 55–6, 78, 83, 92, 100, 102–3, 119, 124, 218, 219, 230, 231, 239, 247, 263, 265, 267–268, 276, 303, 305, 389–90, 519, 567, 569, 585, 590, 592, 602, 610–1 see also climate evaporites 7, 14–5, 36, 94, 127, 171, 250, 261, 264, 276, 292, 389, 425 see also Messinian Salinity Crisis

evapotranspiration 220, 232, 233, 314, 552, 555, 569, 571, 584, 643 see also evaporation; transpiration Evros River, Greece 151 Evrotas River, Greece 328, 330 see also frontispiece extensional tectonics 8–11, 13, 17, 21–3, 36, 195, 385, 451, 473, 480 see also tectonics and landscape development extinction 90, 95, 97, 111–3, 125, 139–40, 142, 145–7, 154–5, 157, 276, 621–2, 626–7, 629–37, 639–40, 642 extinct volcanoes 450 Faraya-Mzaar Mountains, Lebanon 70 farming see agriculture faulting and faults 13, 15, 16, 17, 18, 19, 20, 21, 22, 23, 36, 189, 196, 256, 261, 290, 291, 293, 294, 295, 299, 325, 372, 388, 390, 402, 446, 469, 470, 472, 473, 474, 476, 477, 478, 479–80, 481, 482, 486, 487, 493, 494, 497, 498, 502, 504–505 see also tectonics and landscape development Ferdinandea Island 10 fire see wildfires fish fauna and habitats 143, 154–5, 158–9, 267, 272, 279, 281, 347, 394, 402–3, 621, 627–8, 635–6 see also fisheries and fishing fisheries and fishing 83, 257–60, 266, 272–3, 276–9, 281, 406, 485, 501, 503 flash flooding see storms and floods flint tools 312 see also stone tools flooding see storms and floods flood hazard see storms and floods flood management see storms and floods floodplains 234–5, 242, 245, 246–9, 263, 322, 338, 344–7, 393, 402 see also river systems and environmental change; gravel bed rivers; channel incision; sediment loads and yields fluvial terraces see river terraces; channel incision flysch 171, 246, 294, 298, 300, 325, 333, 404, 455 Fontaine de Vaucluse 16, 288, 296, 298 see also karst foraminifera 47, 50, 51, 54, 56, 92, 103, 122, 123 see also oxygen isotope records forest clearance see deforestation forest communities (past and present) 90, 93–8, 107, 114–7, 123–5, 141–2, 146–8, 150–1, 155–8, 172–3, 177, 180, 183, 187, 203–9, 216, 218,

655

222–4, 237, 246, 249–50, 269–74, 329, 336, 439, 532–3, 541–556, 563–4, 569–72, 574–5, 577, 617–27, 629, 632, 640–4 see also coniferous forest; deciduous forest; rainforest forest fires see wildfires fossil fuels 603, 609 France 16, 37, 74, 77, 78, 84, 86, 109, 114, 116, 124, 141, 146, 159, 160, 171–4, 179–81, 183, 185, 187, 192, 194–5, 206, 211, 213, 217, 222–4, 232–3, 236–7, 249, 255, 263, 273–5, 287–8, 293, 295, 296, 298, 301, 307, 312, 314, 339, 344, 358, 360, 374, 386, 394, 398, 417, 418, 421, 424, 436, 459, 471, 485, 487, 513–4, 516, 521–2, 525, 527, 530, 532–7, 543–4, 552, 554, 563, 566, 583, 587–8, 594–6, 600, 621–6, 628, 631–5, 640–2 see also Corsica Franchthi cave, Greece 288, 312, 402–3, 444, 451, 484 frost and frost action 120, 173, 178, 292, 310, 353, 357, 360–1, 378 see also permafrost; weathering and rock breakdown fuel (for burning and wildfires) 93, 157, 218–9, 220, 543, 546, 548–50, 552–6 see also fossil fuels, ignition FYR Macedonia 10, 11, 233, 256, 257, 258, 274, 275, 279, 367, 479, 587, 594 Galicia, Spain 169, 301, 375 Garda Lake, Italy 257, 262, 280 garrigue 181, 203, 206, 208, 212, 223, 233, 563 gastropods 272, 393, 396, 500 see also molluscs Genoa, Gulf of 38, 39, 74, 75, 77, 78, 80, 390, 497, 515, 537 genetic diversity 109, 112, 113, 117, 125, 126, 142, 143, 152, 160, 216, 623, 624, 628, 638–40, 645 see also genetic divergence genetic divergence 90, 125, 143 see also speciation geoarchaeology 319, 322, 393, 394–5 geochronology see dating methods geophytes 212, 214, 218, 626, 632 geothermal energy 305, 435, 440, 442 Gibraltar 74, 288, 294, 324, 391, 631 Gibraltar arc 11, 12, 13, 15 Gibraltar, Straits of 7, 14, 15, 33, 34, 35, 36, 40–3, 47–48, 77, 78, 94, 119, 144, 211, 386, 389, 390, 401, 406, 532, 601, 610 GISP2 57, 105, 121 see also GRIP

656

Index

glacial and periglacial environments 353–78 Holocene and Little Ice Age records 361–64 interaction with other geomorphological systems 377–8 see also river systems and environmental change modern environments and processes 354–61 Pleistocene records 364–77 Global Positioning Systems (GPS) 8, 10, 11, 195, 472, 473, 476, 480 global cooling 57, 91–2, 95, 119, 122–3, 174, 197, 331, 361 see also Quaternary environmental change global warming and impacts 78–80, 83, 92, 119, 268, 398, 406, 564, 566, 576–7, 599, 610, 624, 632, 643, 644 see also climate; greenhouse gases goats 146, 154, 157, 220–1, 290, 336, 563, 570, 575, 637, 639–40, 642 goethite 176 see also soils gorges 141, 171–2, 194–5, 212–3, 292, 300–302, 305–6, 309–11, 315, 323, 332, 335, 340, 343, 389, 638, 641 Granada, Spain 109, 177, 193, 416, 422, 572 grasslands 93, 107, 116, 146, 151, 203–5, 208–9, 212, 216–7, 223, 550, 556, 570, 578, 617, 626, 631, 634, 643 see also steppe gravel bed rivers 234–5, 240, 242, 245, 297, 298, 305–6, 308, 325, 327, 332, 334, 345, 402 see also river systems and environmental change; sediment loads and yields grazing 184, 203, 204, 212, 214, 216, 217, 218, 220, 224, 225, 237, 239, 249, 272, 273, 438, 546, 548, 549, 553, 554, 563, 567, 571-2, 573, 575-7, 579, 619, 631, 644 see also browsing; herbivory; overgrazing; pastoralism gravel extraction 235, 345–7 Great auk (Pinguinus impennis) 141, 633 Greece 7, 10–11, 19–20, 21–3, 36, 70, 75, 77, 80, 84, 95, 98, 100, 102–5, 107, 109, 111, 113–4, 116–8, 120–2, 123, 124, 125, 127, 143, 146, 150–1, 155, 171, 173, 174, 178,179, 185, 188, 206–7, 212, 213, 215–6, 230, 232, 233, 235, 239, 240, 245, 249, 255, 257–9, 261–2, 266–7, 274, 275, 279–280, 288, 290–1, 296, 303, 306, 311–2,

313, 315, 319–24, 328, 330, 331–3, 337–9, 344, 346–7, 353, 355, 359, 362, 364, 365–72, 376–8, 385, 386, 399–400, 402–3, 405–6, 422, 424, 442–3, 450–7, 459, 469–72, 474, 477, 478–83, 484, 485–7, 497–504, 506–7, 514, 522–3, 530, 533, 534, 541, 543–4, 554, 563, 566, 569, 572, 575, 587, 588, 594, 595, 596, 599–600, 608, 616, 621, 628, 630, 631, 632, 635, 640–1, 642, 644 see also Crete greenhouse effect 592, 599 see also climate; greenhouse gases greenhouse gases 80, 83, 89–90, 92, 529, 599, 607, 609 see also global warming and impacts Greenland ice cores, see GISP2 and GRIP GRIP 57, 329, 330 see also GISP2 groundwater 70, 79, 194, 230–3, 249, 251, 255, 263, 267, 272, 274, 276, 280, 282, 290, 295–8, 301, 303, 305, 307, 309, 312–4, 394, 406, 424, 439, 446, 527, 564, 571, 583, 585–7, 589–94, 621 see also aquifers; water resources Guadalentin River, Spain 243, 249 Guadalfeo River, Spain 193, 249 Guadalope River, Spain 328–30 see also Ebro River basin gully development 5, 18, 20, 22, 23–5, 185–6, 189–90, 246–7, 337, 570, 573, 575 gypsum 16, 20, 26, 171, 185–6, 261, 263–4, 272, 292–3, 297, 305, 423, 425 see also evaporites habitats: aquatic 54, 56, 157, 266, 276, 280–1, 313, 404, 396, 622–43 diversity 149–151, 153, 157, 159, 160, 209, 212, 214–7, 287, 617, 643–6 human impact 146, 158, 217, 274, 276, 313, 346–7, 554, 616–43 terrestrial 107, 120, 125, 139, 141, 145, 147, 149, 156, 209, 618, 622–43 Hadley Cells 38, 122, 126–7, 602 Hadley Centre Regional Climate Model, HadRM3 83–5 haematite 175–6 see also iron oxides halophytes 97 haplotypes 111, 143 harbours see coastal environments hazards see specific hazards heathlands 203–4, 211, 547, 575 heatwaves see climate Heinrich Events 45, 57, 119, 268, 312, 330–1

Hellenic arc 7–9, 12–3, 21, 255, 385–7, 390–1, 393, 473–4, 479, 496–500, 506–10 see also subduction Hellenides, Greece 7, 10, 108, 293 herbaceous vegetation 93, 97–8, 107, 116, 127, 205, 216–7, 219, 548, 626, 633 herbivory 204, 217–8, 556, 575, 633, 637 Hérault River, France 172 Herculaneum see Pompeii Herzegovina see Bosnia hillslope-channel coupling 324, 344 hillslopes 169, 179–97 see also mass movements hillslope terracing see agricultural terraces Holocene 12, 13, 20, 25, 47–9, 55–6, 57–8, 79, 89–90, 99–100, 102–3, 115–7, 122–4, 139, 141–2, 147, 157–8, 178–9, 208, 213, 217, 222–4, 263, 268–72, 303, 307, 310, 312, 315, 320–3, 335–44, 353, 361–4, 366, 369, 378, 388, 390–3, 397–400, 410, 403–4, 423, 441–2, 446, 456, 473, 481–2, 499, 529, 541, 546–8, 564, 575, 617, 634, 638 hydrology see climate; hillslopes; runoff; river regimes; storms and floods Iberian lynx (Lynx pardellus) 156, 642 Iberian Mountains 532, 359–60, 364–5, 375 Iberian Peninsula 11, 21, 36, 37, 71, 74, 75, 98, 102, 111, 120, 139, 141, 142, 144, 158, 159, 206, 212, 222, 225, 233, 239, 254, 261, 324, 329, 331, 359–60, 364–5, 375, 376, 515, 521, 546, 547, 568, 584, 596, 617 see also Spain; Portugal Ibex (Capra ibex) 145, 156, 633–4 ice sheets 37, 47, 90–2, 95, 98, 102, 104, 107, 113, 122, 125, 262, 331, 373–4, 376, 392 402, 435 see also glacial and periglacial environments; ice caps ice caps 47, 299, 354–5, 364, 367, 375 ice cores see GISP2, GRIP, Antarctic ice cores ice rafting see Heinrich Events ignition 218, 220, 543, 546–9, 553–4, 556, 609 see also fuel; wildfires incision see channel incision; gully development; river terraces India 73, 91, 602–4 Indian monsoon 39, 46, 55, 56, 73, 74–5, 602–3 Indian Ocean 7, 14, 94, 509–10, 603

Index interglacials 7, 45, 46, 47, 51, 55, 56, 57, 58, 90, 92, 94, 97–100, 101, 102, 104–7, 114–9, 122–5, 125, 127, 140, 141, 203, 212, 215, 216, 268, 312, 319, 320, 321, 325, 327, 334, 377, 390, 391, 546 see also Holocene; Quaternary environmental change; Last Interglacial insolation 45–6, 54, 55, 56, 58, 98, 99, 100, 101, 102, 106, 115, 116, 117, 118, 119, 123, 125, 127, 175, 222 see also climate; Milankovitch cycles Inter-governmental Panel on Climate Change (IPCC) 80, 83, 527 interrill erosion 183, 185 Intertropical Convergence Zone (ITCZ) 38, 99, 175, 176, 178, 602, 603, 605 invasion and invasive species 146, 147, 157, 207, 209, 214, 220, 222, 616, 626–8, 629, 630, 633, 644 invertebrates 148–149 Ioannina Lake basin, Greece 90, 100, 101, 103, 104, 105, 107–8, 109, 115, 117, 120, 121, 124, 125, 216, 259, 261, 262, 362, 365, 369, 377, 572 Ionian Islands 401, 471, 477, 485 Ionian plate 21, 22 Ionian Sea 10, 15, 22, 41, 42, 44, 74, 76, 77, 78, 107, 385, 387, 388, 400, 404, 406, 471, 472, 474, 479, 480, 486, 505, 508 Iran 255, 257, 259, 261, 274, 486, 602 Iraq 75, 86, 255 Iron Age 158, 223, 271, 339 iron oxides 175, 176, 178 see also terra rossa; soils irrigation 40, 70, 79, 83, 232, 234, 239–40, 249, 250, 257, 260, 269, 272, 274, 276, 277, 278, 280–1, 345, 347, 564, 571, 584, 585, 589–90, 591, 592, 596, 635 Israel 18, 19, 56, 74, 80, 97, 99, 103, 104, 116, 119, 171, 173, 175, 176, 178, 182, 184, 206, 207, 208, 209, 215, 217, 220, 232, 233, 236, 244, 247, 255, 257–60, 274, 275, 276, 277, 280, 288, 312, 313, 315, 386, 389, 394, 405, 418, 421, 422, 423, 424, 471, 477, 480, 496, 498, 500, 504, 505, 535, 541, 563, 587, 588, 589, 593, 594, 595, 596, 616, 621, 625, 631, 636, 640, 642, 644 Istanbul 278, 279, 478, 504 see also Constantinople Italy 7, 8, 10, 11, 16–7, 19, 20, 21, 36, 38, 69, 75, 77–80, 83, 85–6, 92, 97,

102, 109, 113, 116–117, 119–20, 123–4, 141, 157, 171, 173, 177–8, 185, 188–9, 192–3, 195–6, 206, 222, 232–4, 236, 255, 257–62, 267–9, 270, 274–8, 280, 287–8, 292, 295, 298, 303, 312–4, 319, 331, 339, 345, 357–8, 375–6, 385–6, 391, 395, 400, 406, 421–3, 435–7, 440–1, 442–50, 457, 459, 469–71, 474, 477, 478, 480, 484, 487, 495–7, 506–10, 513–4, 521–2, 525, 527, 529–30, 532–7, 541, 543–4, 554, 565–6, 569, 572, 587–8, 594–6, 600, 621–2, 630–1, 635, 640, 642 see also Appenines, Calabria; Sicily; Sardinia ITCZ see Intertropical Convergence Zone jet stream 38, 73, 122, 531, 537, 602–4 Jerusalem 312, 420, 535 Jordan 206, 275, 631, 642 Jordan River 236, 265, 276, 277, 280, 483, 485 Jordan Valley and rift 56, 255, 256, 257, 261, 265, 266, 268, 272, 276, 295, 321, 324, 418, 469, 471, 480, 504, 505 Julian Alps 331, 354, 357, 362, 378 Jura Mountains 11 Jurassic Period 94, 127, 293, 294, 295 karst 7, 16, 26, 287–315 see also tufa; travertines; speleothems caves 298–300, desert karst 303, 305 early history of karst research 290 fauna 156, 157 glacio-karst 291, 314, 353, 364, 367, 377 see also glacial and periglacial environments habitats 150 hydrology 232, 295–8, lakes 257–60, 261, 262, 264, 266, 272, 273, landforms 290–310 landscapes 255, 290–5, 300–307 records of environmental change 310–13 threats and conservation 312–4 vadose and phreatic systems 295–300 weathering 171, 174, 292 Kastritsa cave, Greece 103, 104, 105, 108 katabatic air 37, 77, 78 see also climate Konya basin, Turkey 258, 260, 261, 268, 274, 275, 416, 423, 621 Kopais basin, Greece 90, 105, 109, 120, 121 Köppen classification of climate 69 see also climate

657

Korana River, Croatia 288, 306 Krka, River, Croatia 306 Lago Mare 14, 36 see also Mediterranean Sea lagoons 23, 40, 83, 153, 255, 258, 260, 263, 272–3, 278, 279, 280, 293, 386, 393, 402, 406 see also lakes and wetlands lake levels 47, 56, 57, 102–5, 119, 123, 268–70, 278, 280, 340, 342 lakes and wetlands 255–82 biology 266–7 hydro-chemistry 263–66 long records of change 267–72 origin and morphology 256–63 recent trends 272–4 threats and conservation 274–81 lake sediments 23, 116, 178, 262, 272, 280–281, 420, 444, 546, 563 laminated sediments 37, 48, 51, 102, 123, 258, 267–8, 447 land abandonment 212, 214, 217, 224, 271, 485, 549, 553, 555, 566, 571, 572–5, 618 land degradation 563–79 see also habitats; soils; biodiversity and conservation contemporary perceptions 565–7 history 563–4 processes and precursors 567–77 prospects for improvement 577–9 landslides, 20–21, 24–6, 73, 79, 86, 189–90, 192–6, 334, 378, 437, 439, 441, 445, 448–9, 471, 482–5, 493, 497, 501–3, 505, 510, 527, 534–5, 565 see also mass movements Languedoc-Roussillon, France 516, 522, 525, 536 La Niña 73 Last Glacial Maximum (LGM) 45, 47–9, 55, 102–4, 106, 107–12, 114, 268, 375–6, 393 Late-glacial 20, 102, 116, 127, 141, 171, 268, 321, 327, 333, 367, 370, 373, 375 Last Interglacial 7, 99, 100, 101, 115–7, 122, 123–4, 325, 327, 390, 391, 546 lava 175, 257, 335, 355, 437, 438–42, 444, 446–51, 453–5, 458–9 see also volcanoes and volcanism; magma and magmatism Lebanon 70, 108, 149, 207, 209, 215, 255, 275, 295, 299, 364, 386, 418, 480, 495, 504, 505, 542, 587, 594, 595, 616, 631, 632, 640, 644 Lesvos, Greece 174–5, 216, 575 Levant 40, 78, 99, 145–6, 169, 232–3, 239, 255, 261, 269, 271, 287, 293,

658

Index

295, 326, 385, 386, 392–3, 400, 406, 415, 418, 424, 495, 504–5, 563–4, 583 Levante, Spain 78, 516, 521, 532, 536–7 Levante wind 77, 78, 532, 537 see also climate Levantine Sea 33, 34, 36, 39, 41, 42, 44, 385–7, 388, 496, 504–5, 508, 510 levées 20, 454 Libya 10, 55, 58, 78, 145, 156, 169, 215, 232, 241, 274, 275, 287, 288, 295, 301, 305, 315, 319, 324, 325–9, 330–1, 386, 416–8, 422, 424, 470, 477, 480, 485, 508, 583, 586–9, 592–6, 630, 631, 640 see also; Cyrenaica; Tripolitania lichens 173, 244, 422 lichenometry 340, 343, 344, 362, 364 Liguria 496–7, 508, 514, 516, 521, 525, 536 Ligurian Alps 108, 215, 616, 644 Ligurian Sea 38, 39, 75, 387 limestone 7, 16, 25–6, 171, 174–5, 178, 179, 233, 255, 261, 287–315, 321, 325, 327, 332–3, 353, 365, 367, 378, 385, 394, 396, 399, 416, 450, 455, 476, 586, 626, 636, 641 see also karst; calcium carbonate Lions, Gulf of 35, 37, 39–40, 42–3, 103, 197, 401, 405, 424 Lisan, Lake 103, 105, 119, 257, 268 see also Dead Sea Lisbon 69, 71, 480 Little Ice Age (LIA) 79, 249, 272, 321, 340–44, 348, 353, 357, 361–4, 404 littoral zone 250, 274, 276, 307, 329, 331, 385, 386–90, 392–407, 415, 439, 523, 530 see also coastal environments Llobregat River, Spain 516, 523–7, 530, 535 long profiles of river channels 325, 335 see also river systems and environmental change; tectonics and landscape development Lost Eden hypothesis 214–5, 566, 615 luminescence dating 321, 323, 329, 422 optically stimulated luminescence dating (OSL) 325, 328, 330, 334 thermoluminescence (TL) dating 332, 375 Lyell, Charles 398, 450 Macedonia, Greece 10, 151, 276, 338, 472, 477, 485, 621 Maggiore Lake, Italy 255, 257, 262, 273, 278, 280

Maghreb 10, 11, 16, 107, 169, 233, 239, 245–7, 255, 325, 331, 421, 565, 566, 571, 583, 588, 619, 630 magma and magmatism 10, 13, 14, 436, 439, 441–51, 452, 454–8 see also volcanoes and volcanism Majorca 240, 299, 309, 310, 554, 586 see also Balearic Islands Malta 10, 14, 70, 275, 386, 401, 533, 587–9, 594, 596, 621, 631, 634, 638, 640 Mammal fauna 93, 141, 143–4, 145–7, 154–8, 214, 218, 626, 628–9, 633–4, 638, 642 mammoths 145, 146, 633 maquis 151, 173, 188, 203–4, 206–7, 209, 211, 212, 214, 233–4, 337 marine environment see Mediterranean Sea marine gateways 5, 7, 14–5, 91, 94, 389 see also Mediterranean Sea marine terraces 5, 7, 22, 390 see also coastal environments Maritime Alps 108, 215, 354, 358, 362, 373, 375–6, 378, 616, 621, 626, 644 marls 7, 19, 26, 173, 178, 180, 183, 185, 189, 246, 248, 250, 263, 272, 293, 325, 333 Marmara, Sea of 274, 390, 495–6, 504, 505, 506, 508, 510 see also Black Sea Marseille 16, 74, 77, 141, 395, 398, 633 marshes 23, 151, 213, 255–6, 260–1, 263, 272–3, 274, 276–7, 280–1, 386, 402, 536, 590, 621, 623, 628, 636, 642 see also lakes and wetlands mass movements 5, 7, 17–8, 20–1, 26, 169, 188–91, 197, 311, 514 see also debris flows; landslides; pyroclastic density currents and deposits Massif Central 77, 109, 113–5, 436, 532 mattoral 147–8, 150, 152, 156, 180, 183–4, 190, 203, 206, 249, 618, 625, 627, 629 MEDALUS 249, 577 Medieval Period 322, 335, 338, 339, 546 see also Medieval Warm Period Medieval Warm Period 79, 321, 341 Mediterranean Sea: present and past 33–58 basin topography and water circulation 33–6 centennial- to millennial-scale variability 57–8 evolution of the basin 36–7 glacial cycles 47–8

modern climate and oceanography 37–40 monsoon maxima and sapropels 48–57 Quaternary change 44–58 water circulation and deep water formation 40–44 Megara basin, Greece 22–3 Mesolithic Period 158, 310, 314, 338, 339, 401, 403, 406, 451 Mesozoic Era 10, 33, 36, 94, 255, 261, 295, 307, 324, 445, 455 Messina, Italy 471, 478, 497 Messina, Straits of 400, 495–6, 497, 505, 508–10 Messinian Salinity Crisis 5, 7, 14–6, 19, 23, 26, 36, 94–5, 127, 144, 151, 171, 195, 261, 287, 292, 296, 307, 388–90 see also tectonics and landscape development; evaporites metal mining 345–6 metal pollution 274, 277, 280, 282, 345–6, 439, 442, 450, 450–1, 638 Methana, Greece 13, 450–1 methane 80, 124, 406, 599, 603, 607 see also greenhouse gases microclimate 108, 140, 171, 173, 179, 212, 215, 217, 357 microplates 139, 144, 293, 295, 324 see also tectonics and landscape development Middle East 39, 74, 75, 77, 83, 84, 86, 146, 147, 204, 324, 418, 420, 422, 471, 478, 608, 610, 630, 640 Mid-Pleistocene Transition (MPT) 37, 106, 114 migration (flora and fauna) 90, 94, 95, 107, 109, 112, 113, 120, 140, 142, 144–6, 151, 152, 154, 155, 156, 212, 214, 234, 266, 274, 276, 281, 624, 635, 637–8 see also invasion and invasive species Milankovitch cycles 45, 58, 89, 90, 124, see also sub-Milankovitch climate variability eccentricity 37, 45–6, 54, 91, 92, 98–9 obliquity 37, 45–6, 91–2, 98–9, 104, 113, 117, 119 precession 45–6, 54, 97, 98–102, 104, 117–9, 125, 127 Minoan Civilization 214, 436, 441, 443, 452, 454, 460, 493, 499–501, 505 Miocene 5, 10, 11, 13, 14, 21, 92, 93, 94, 95, 97, 98, 126, 127, 139, 142, 144, 145, 178, 216, 256, 261, 305, 307, 455 see also Messinian Salinity Crisis

Index Mistral 35, 37, 40, 43, 75, 77, 78, 532 see also climate; storms and floods molluscs 266, 267, 396, 402, 403 see also gastropods monsoon 36, 37, 39, 40, 45, 46, 48–58, 73, 74, 75, 93, 98, 99, 100, 102, 104, 116, 602, 603, 604 see also sapropels; climate Montenegro 257, 263, 264, 266, 275, 288, 289, 290, 291, 292, 301, 306, 308, 314, 353, 354, 356, 357, 358, 367, 373, 374, 378, 495, 497, 508, 587, 588, 594 Monticchio, Lago Grande di 90, 109, 115, 120, 261, 268, 270, 572 moraines 329, 353, 357, 362–76, 377, 378 sea also glacial and periglacial environments; till Morocco 10, 11, 14, 15, 19, 21, 36, 39, 42, 43, 58, 94, 108, 109, 116, 123, 150, 176, 205, 206, 212, 232, 233, 234, 236, 241, 246, 247, 248, 259, 261, 269, 272, 273, 275, 287, 288, 303, 328, 330, 354, 360, 386, 416, 417, 424, 480, 536, 541, 542, 543, 564, 565, 587, 590, 591, 594, 595, 596, 618, 620, 621, 631, 632, 636, 640, 642, 643 see also Atlas Mountains mudflows 189, 249, 437, 439 see also mass movements Murcia, Spain 178, 189, 240, 242, 243, 245, 249, 250, 514, 530, 573, 576 Nearctic region 142, Near East 17, 78, 80, 89, 100, 104, 107, 116, 147, 209, 213, 222, 223, 224, 319, 364–5, 423, 566, 583, 584, 588, 593, 596, 618, 622, 632 Negev Desert 19, 247, 248, 416, 422, 423, 425 Neogene 23, 24, 49, 91, 127, 145, 293, 481 Neolithic Period 146, 147, 154, 158, 218, 223, 315, 322, 335, 338, 339, 398, 402, 403, 442, 456, 563 Nestos River, Greece, 151 Nile River 15, 16, 17, 18, 36, 39, 40, 54, 55, 58, 83, 97, 99, 100, 116, 176, 230, 232, 233, 258, 263, 272, 277, 386, 388, 389, 402, 405, 420, 422, 530, 584, 585, 588, 590, 594, 596 nitrous oxides 80, 599, 600, 605, 607 see also air quality nivation 174, 302, 353, 357, 357, 364, 365, 378 see also glacial and periglacial environments North Africa 10, 11, 33, 37, 39, 55, 58, 74, 75, 77, 78, 84, 89, 99, 104,

107, 116, 142, 144, 145, 146, 147, 148, 149, 150, 159, 204, 208, 209, 223, 224, 232, 239, 247, 255, 261, 262, 263, 276, 281, 293, 303, 305, 306, 310, 315, 319, 324, 325–9, 331, 385, 386, 392, 400, 401, 415, 416, 418–25, 496, 508, 531, 532, 536, 543, 548, 549, 553, 554, 564, 566, 583, 600, 604, 610, 618, 621, 630, 634, 636, 640, 642, 643 see also Algeria; Egypt; Libya; Maghreb; Morocco; Tunisia North Atlantic Deep Water (NADW) 35, 94, 119, 122 North Atlantic Ocean 7, 8, 12, 14–5, 23, 33, 35, 36–8, 39, 40–2, 45, 47, 57–8, 69–70, 73, 74, 80, 91, 94, 98, 99, 104, 119–20, 122–5, 141, 147, 230, 268, 269, 312, 329, 331, 340, 341, 342, 346, 389, 390, 406, 424, 469, 472, 509, 515, 518, 521, 534, 601, 602, 605, 608, 610 see also Heinrich Events; North Atlantic Deep Water; North Atlantic Oscillation North Atlantic Oscillation, NAO 38, 73, 80, 340, 415, 574, 601, 608, 610 Northern Hemisphere 37, 38, 45, 46, 57, 58, 69, 70, 91, 92, 96, 122, 142, 146, 605, 633 oases 97, 288, 305, 571 obliquity see Milankovitch cycles obsidian 447, 451, 456 Older Fill 321–3 see also river systems and environmental change Oligocene 8, 10, 11, 16, 21, 91–3, 95–8, 126, 145, 216 olive 77, 151, 160, 184, 206–7, 210, 214–6, 220, 271–2, 298, 336–7, 572, 633, 639 Olympia, Greece 319–21, 338, 346, 483 Olympus, Mount, Greece 108, 354–5, 366–7, 372 ophiolite 333, 353 orbital forcing see Milankovitch cycles orbital tuning 46, 94, 369, 436 see also Milankovitch cycles; dating methods orogeny 5–6, 10–1, 197, 255, 331, 334, 473 see also tectonics and landscape development orographic influence on climate see climate; storms and floods Ottoman Period 271, 638 outwash sediments 372 see also glacial and periglacial environments overbank flows 247, 527, 529 see also storms and floods; floodplains overbank sediments 247–8, 333, 340–1 see also floodplains

659

overgrazing 158, 208, 272, 405, 564, 570, 571, 577, 615, 618, 632 see also grazing; pastoralism overkill hypothesis 146, 634 overland flow 182–3, 185, 239, 244, 247, 249, 575 see also runoff; hillslopes oxygen isotope records: ice cores 57, 120, 328, 330, marine 36, 46, 47, 53, 54, 55, 56, 58, 91, 92, 95, 98, 105, 106, 113, 114, 118, 120, 121, 122, 124, 268, 312, 329, 365, 369, 376, 390 lake 123, 124, 265, 268, 269, 272 speleothems 312, 313 ozone 122, 217, 599, 604–6, 607, 608 see also air quality; TOMS Pamukkale, Turkey 288, 306, 308, 314, 315 see also travertine; tufa Palaearctic 141–2, 146, 147, 149, 155 palaeofloods and palaeohydrology 272, 339–45 see also river systems and environmental change Palaeolithic Period 157, 310–1, 314, 319–321, 333, 347, 385, 394–5, 401–3, 406, 443–4, 563, 633 Palaeolithic art 314, 395, 633 palaeoshorelines 266, 267–78, 390–403 see also sea-level change; lake levels; marine terraces palaeosols 93, 98, 127, 174, 177–8, 421–2, 424 see also soils Palestine 160, 232, 255, 423, 496, 500, 505, 564, 587 palynology see pollen and pollen records Panama, Isthmus of 91 panarchy 565, 573, 577–8 Pantelleria 10, 13, 263, 435, 436, 449–50 see also volcanoes and volcanism paraglacial 378 see also glacial and periglacial environments parent material see soils pastoralism 179, 405, 423, 564, 575, 577, 616–7, 625 see also grazing passerines 152, 624 see also bird fauna Passer Valley, Italy 193, 195 peat 116, 280, 393, 443, 571 pedogenesis see soils Peloponnese, Greece 239, 321, 323, 337–8, 403, 474–5, 498, 503, 544, 632, 640 see also Corinth; Sparta Pergusa Lake, Sicily 259, 263–4, 268 periglacial processes see glacial and periglacial environments perihelion 45, 46, 54 permafrost 357–60, 378 see also glacial and periglacial environments photosynthesis 93, 204, 224, 552

660

Index

Piave River, Italy 194, 345 Picos de Europa, Spain 288, 295, 298–9, 303, 359, 364, 375, 378 see also Cantabrian Mountains Pindus Mountains, Greece 19, 21, 69, 103, 107–8, 113, 120, 155, 212, 301, 303, 314, 331, 353–4, 355, 357, 364–7, 369, 377–8 Piva River (Sinjac), Montenegro 298, 367 plate tectonics see tectonics and landscape development playa lakes 258, 260, 261–3, 425 see also saline lakes and wetlands Pleistocene see Mid-Pleistocene Transition (MPT); Quaternary environmental change; Last Glacial Maximum (LGM) Pliny the Younger 435, 444, 632 Plinian and subplinian eruptions 435, 442, 444–6 Pliocene 10, 13, 14–16, 24, 36, 49, 91–5, 97–8, 106, 112, 114, 116–8, 125, 139, 142, 145–7, 177, 305, 315, 333, 335, 404, 451, 546 ploughing 182, 222, 239, 249, 336 see also tillage Po River, Italy 39, 44, 389 Po delta 18, 151, 263, 388, 404, 406 pollen and pollen records 47, 56–7, 89–128, 174, 177, 206, 209, 211, 214, 216, 222, 268–72, 328–30, 339, 365, 376–7, 546, 571–2 pollution, see air quality; water quality Pompeii and Herculaneum 444–6 see also volcanoes and volcanism Pontic Mountains, Turkey, 108, 149, 354, 355, 364, 378 Portugal 124, 178, 180, 185, 206, 211, 216–7, 219, 223, 233, 275, 280, 360, 375, 541, 543–5, 554, 566–7, 599–600, 617, 628, 631, 640 Portugese margin 120–1, 124 potassium-argon dating 335, 355 pottery 321, 456, 476 see also artefacts Precambrian 324, 325 precession see Milankovitch cycles precipitation see climate progradation see coastal environments; sea-level change Provence 249, 393, 470, 471, 485, 516, 517, 535 Provençal basin 7, 9, 10, 15 Pyrenees 10, 19, 70, 79, 80, 84, 108, 109, 141, 144, 149, 171, 178, 212, 293, 298, 301, 314, 325, 331, 354, 358–9, 360, 362–4, 373–5, 376, 377, 378, 404, 417, 515, 516, 521, 522, 523, 524, 532, 573

pyroclastic density currents and deposits 437, 438, 441, 442, 444, 445, 446, 447, 452, 455, 456, 457, 458, 500, 505 Qattara depression, Egypt 416, 425 Quaternary environmental change: see also Last Glacial Maximum (LGM); Heinrich Events; Dansgaard-Oeschger events; interglacials; Last Interglacial aeolian processes 415–25 climate and vegetation change 98–128 coastal environments 388–402 glacial and periglacial environments 361–77 history of vertebrate fauna 139–160 karst dynamics 310–13 marine environments 44–58 river systems and environmental change 325–40 seismicity 469–84 tectonics and landscape development 5, 7, 15, 16–26, volcanism 436, 441–57 radiocarbon dating 104, 141, 312, 321–3, 332, 335, 338, 340–2, 362, 369,374,375–6,400,444,454,505 AMS dating 104, 323, 332, 340–2 databases 340–2 raindrop impact 179–80, 183–4, 219, 572 see also soils rainfall see climate rainforest 94, 95 raised beaches 23, 481 see also marine terraces; palaeoshorelines sea-level change ramblas 24, 26, 186, 242–3, 245, 249 see also wadis Ramsar convention and designated sites 256–60, 274–5, 277, 640 raptors 152, 637, 642 see also bird fauna reafforestation 237, 271 see also afforestation recharge see groundwater red beds 321–2 see also terra rossa Red deer (Cervus elaphus) 145, 642 Red Sea 10, 33, 47, 56, 99–100, 144, 261, 390, 418, 472 refugia 90, 107–14, 125–6, 140–3, 212, 216, 621, 643, 645–6 relictual taxa 113 Renaissance 313 reservoirs see also dams; water resources evaporation losses 590 sedimentation 193, 239–42, 244, 247, 345, 571, 591–2 water storage 79, 194, 196, 239–43, 277, 279, 345, 585, 589–92, 642

Rhodes 18, 33, 42, 471, 473, 478, 482, 495, 498–9 Rhône River and valley, France 17, 35, 37, 39, 77, 86, 151, 230, 233, 236, 344, 404, 527, 584, 585, 586, 621 Rhône River delta and sub-marine fan 18, 20, 263, 346, 388, 393, 402, 404, 405, Rhône River palaeovalley 15, 16, 195, 389 RICAMARE 644 Rifian corridor 7, 14, 15 see also Rif Mountains; Betic-Rif region Rif Mountains 10, 11, 19, 24, 150, 215, 616, 644 see also Betic-Rif region rifting 10, 195, 261, 324, 385, 436, 472 see also rift valleys rift valleys 261, see also rifting; Jordan Valley and rift rillenkarren 292, 310 see also karst; weathering and rock breakdown rills and rill erosion 185–186, 187, 239, 249 see also hillslopes riparian habitats 95, 250, 346, 585, 621, 627, 628, 629, 632 see also wetlands river capture 7, 16, 21, 23, 24–6, 333–4, 335 river channels 150, 189, 235, 248, 319, 321, 325, 336, 337, 340, 344–7, 480, 481, 483, 522, 523, 529, 537, 584 river regimes 47, 55, 229–34, 235, 237, 239, 240, 243, 248, 250, 315, 321, 323, 344, 523, 536, 584–6 see also climate river systems and environmental change 319–48 see also sediment loads and yields Holocene records 335–340 Little Ice Age 349–344 models of change 321–4 Pleistocene records 325–35 see also river capture; river terraces recent human impacts 344–7 tectonic setting for 324–35 see also river capture river terraces 15, 18, 321, 322, 325, 327, 328, 329, 330, 333, 334, 335, 336, 344 rock breakdown see bedrock erosion and weathering rockshelters 108, 287, 297, 310, 321 see also rockshelter sediments, caves and cave deposits rockshelter sediments 174, 297, 298, 310–2, 313, 333, 378, 484 Roman Period 160, 179, 239, 240, 247, 263, 271, 272, 276, 281, 305, 306, 321, 322, 327, 335, 336, 339, 395,

Index 398, 399, 400, 404, 405, 442, 445, 446, 451, 476, 480, 485, 529, 546, 563, 564, 617, 628, 632 see also Classical Period Roman triad 158, 616 Romania 11, 143, 275, 550, 609 Roman volcanic province 436, 442 see also volcanoes and volcanism Rome 69, 71, 236, 276, 306, 404, 437, 442, 621 see also Roman Period runoff 26, 39, 54, 55, 56, 57, 99, 102, 104, 151, 158, 178, 180, 182–7, 214, 229, 232, 233, 235, 237, 239, 240, 244–5, 249, 251, 272, 280, 314, 319, 331, 337, 477, 536, 552, 553, 566, 569, 570, 572, 573, 574, 575, 579, 583, 584, 587, 591, 618, 626, 643 see also hillslopes; climate; storms and floods; river regimes; Mediterranean Sea Sahara Desert 55, 58, 69, 75, 77, 97, 99, 116, 141–2, 145, 152, 255, 263, 360, 415–22, 424–5, 531, 571, 577, 583–4, 586, 592, 637 Saharan dust see aeolian processes and landforms Sahel 55, 58, 74,116, 602, 610 saline lakes and wetlands 94, 123, 213, 250, 258, 260–7, 269, 274, 276–8, 280, 346, 389, 394, 425, 485, 571, 592 saline soils 211, 249 see also salinization salinization 70, 179, 272, 274, 276, 346, 571 salt deposits see evaporites salt weathering 425 Santorini (Thera) 13, 269–70, 400, 435–6, 441, 443, 451–6, 459–60, 485, 493, 495–6, 499–502, 505 sapropels 33, 36, 46, 47, 48–57, 98–102, 104, 116, 123, 128, 267, 269, 312–3 see also Mediterranean Sea Sardinia 7, 10, 42, 232, 263, 324, 329, 393, 394, 417, 420–1, 424, 514, 516, 521, 547, 571, 631 satellite imagery and data 76, 108, 406, 417, 476, 534, 544–5, 550, 553, 555, 565, 576, 600–1, 605, 608–9 see also TOMS savanna and savanna-like habitats 58, 93, 99, 109, 116, 146, 149, 203–4, 206, 246, 577, 615 sclerophyllous vegetation 90, 97–8, 100–1, 109, 115–7, 126–8, 173, 203–7, 212, 222–4, 249, 546, 623

scree 171–2, 178, 205, 212–3, 303, 321, 322, 324, 370, 378, 641, 643 see also talus sea-level change 7, 14–6, 18, 21–3, 36, 47–8, 80, 83, 91–2, 94, 103, 141, 195, 213, 263, 303, 305, 307, 312, 323, 338, 385, 388, 390–406, 423, 425, 482–3, 497, 531, 643 sebhkas 261–2, 405, 424–5 sediment budgets 248, 298, 300, 346, 347, 386 sediment supply 54, 235, 319, 323, 331–3, 335, 340, 344, 353, 378, 402, 483 see also hillslope-channel coupling; sediment budgets sediment loads and yields see also reservoirs suspended sediment 17, 23, 26, 40, 54, 58, 186, 195, 214, 237, 242, 245–8, 297, 420, 591 bed load sediment 17, 23, 26, 195, 247–8, 297, 569–71 Segre River, Spain 516, 530 seismicity 5, 7, 10, 11–13, 15–6, 18, 20–23, 26, 139, 189, 324–5, 329, 331, 387, 399–401, 404, 439, 441–2, 444–52, 454–5, 457, 459, 469–87, 493–4, 497, 500–2, 504–9 see also earthquakes; tsunamis Serbia 11, 45, 367, 587–8, 594 Sicily 10, 11, 12, 13, 17, 20, 77, 97, 113, 117, 144, 169, 206, 222, 247, 259, 263, 264, 268, 325, 335, 336, 357, 390, 395, 401, 415, 418, 435, 446, 449–50, 472, 476, 495, 496, 497, 498, 508, 514, 571, 631, 632, 634 see also Etna Sicily–Cap Bon line 139 Sicily, Strait of 10, 33, 34, 35, 41, 42, 52, 75, 401 Sierra Nevada, Spain 108, 174, 212, 249, 255, 354, 359–60, 364, 378 Sinai 149, 324, 326, 386, 416, 418, 422–3, 425 sinkholes 303, 305, 484 see also karst Sirocco wind 74, 77, 415, 418, 513, 531, 537 see also climate; storms and floods Sirte basin 10, 388, slackwater deposits 311, 333, 340–1, 348 see also palaeofloods slope deposits see scree; hillslopes slope failure see mass movements Slovenia 11, 109, 143, 169, 293, 298, 314, 331, 357, 362, 377, 525, 533, 541, 587–8, 594, 636, 640 snowline 354–5, 357, 364, 375–6 see also glacial and periglacial environments, Equilibrium Line Altitude (ELA)

661

snowmelt 91, 192, 439, 527 soils see also terra rosse aggregates 173, 185 carbon 172, chronosequence 174 conservation 245, 565, 578–9 erosion 179–188, 190–192, 197, 209, 214, 219, 237, 247, 249, 271–2, 336–7, 552–3, 563, 566, 574, 576–7, 579, 618, 626, 638, 643 forming processes 173–4, 176–7 moisture 102, 180, 206, 217, 232, 240, 247, 583 nutrients 173, 566 organic matter, 172–3, 220, 301, 566, 572, 574 parent material 169, 171, 174–175, 178 profiles 172, 176, 178–9, 290, 332 resources 178 solifluction 179, 355, 357, 360, 376 see also glacial and periglacial environments; hillslopes solute loads 176, 246, 263, 265, 290, 297, 301, 305, 442 see also dissolution; sediment loads and yields Sorbas basin, Spain 14, 24–6, 171, 245, 288, 293, 333–4 Soreq Cave, Israel 56, 90, 102–3, 123, 288, 312–3 see also speleothems South Africa 69, 157, 222, 541, 546–7, 552, 575, 578, 615, 627, 628 Spain 5, 7, 10–11, 13, 18–20, 23–26, 41, 70, 75, 77–9, 80, 84, 86, 94, 108–9, 116, 123, 146, 156, 169, 171, 173, 174, 177–8, 180, 182, 183, 185–93, 205, 206, 209, 211–13, 217, 232, 237, 240–5, 248–50, 258–65, 266, 267, 269–70, 274–6, 277–79, 280, 288, 293, 298–9, 301, 309, 314, 319, 324, 328–31, 333–5, 339–42, 344–6, 353, 359–60, 362–4, 373–6, 388, 390, 401, 405–6, 415–6, 420–1, 423–5, 436, 477, 513–4, 516–7, 519, 521–2, 528–38, 541, 543–44, 546–9, 552, 554–5, 564, 566–9, 572, 574, 576, 586, 587, 590, 594, 605, 623, 625, 631, 640 Sparta, Greece 471, 485, 502 see also frontispiece speciation 90, 125–6, 139–40, 148, 160, 216, 624 see also allopatric speciation; genetic divergence species diversity 139, 143, 217, 220, 396, 615–6 see also biodiversity

662

Index

speleothems 56, 101–2, 103, 119, 123, 175, 297, 298, 299, 305, 312–3, 315, 319, 393–4, 420 steppe 37, 93, 97, 107, 115, 118, 122, 141, 147, 203, 208–10, 212, 223, 246, 268, 329–30, 564, 618–9, 634, 643 stone-walled terraces see agricultural terraces stone tools 311, 319, 321, 444 see also Palaeolithic Period storms and floods 513–38 see also catastrophic flooding classification of floods and storms 523–27 climate variability and flooding 527–31 convective systems and storms 518–23 cyclones and cyclogenesis 38, 39, 58, 71, 74–77, 126, 513, 515, 517, 518–9, 521, 522, 531, 532, 534, 535, 536, 537, 600, 605 see also anticyclones hailstorm events 534 hazard mitigation 536–7 heavy snowfall 534–5 hurricanes 534 impact of extreme events 535–6 tornadoes 532–4 wind storms 531–2 strontium isotopes 102, 420 subduction 5, 7–17, 21, 36, 293–4, 385, 387, 389–91, 435, 450, 456, 474, 478, 498 see also back-arc basins; Calabrian arc; Hellenic arc; tectonics and landscape development sub-Milankovitch climate variability 45, 58 see also Dansgaard-Oeschger events; Heinrich Events; Little Ice Age; Medieval Warm Period sub-tropical high pressure 37, 38–9, 69, 73, 122, 126–7, 514 sub-tropical jet 73, 122, 537, 602, 604 succession, ecological 101, 112, 115–20, 127, 151, 156, 217, 223, 549, 556, 565, 578, 626 Super-Sauze earthflow, French Alps 191, 195 suspended sediment see sediment loads and yields Syria 75, 86, 109, 116, 142, 146, 209, 215, 255, 259, 271, 275, 298, 386, 418, 471, 483, 495, 505, 587, 594–6, 616, 631, 634, 636, 640, 642, 644 Tabernas basin, Spain 18–9, 187 talus 302, 309, 359, 366, 378 see also scree Tara River, Montenegro 367

Taurus Mountains, Turkey, 149, 212, 255, 301, 354–5, 364–5, 378 tectonics and landscape development 5–32 see also earthquakes and seismicity; volcanoes and volcanism; coastal environments global setting and history 5–7 , 14 see also Messinian Salinity Crisis geodynamics of the Mediterranean 7–14 see also earthquakes and seismicity; volcanoes and volcanies temperature see climate, climate change Tenaghi Philippon, Greece 107, 109, 113–5, 118–9, 120–1, 174, 261 tephra 128, 269–70, 311–2, 425, 436–9, 441–8, 450–2, 454–8, 499–500 see also volcanoes and volcanism Ter River, Spain 516, 528, 530 terraces see agricultural terraces; marine terraces; river terraces terra rossa 175, 178, 298, 305, 309, 315, 325, 327, 415–6 see also soils Tertiary Period 10, 20, 94–8, 112–3, 116, 125, 126–8, 139, 142, 178, 212, 255, 305, 425, 631 Tethys 5, 7–10, 14, 15, 36, 94, 95, 97, 127, 144, 145, 171, 255, 261, 293, 294, 385, 386, 387, 472 see also marine gateways Thera see Santorini Thermaikos Gulf, Greece 347 thunderstorms see climate; storms and floods Tibetan high 602, 603 Tibetan Plateau 37, 91, 197 Tigalmamine Lake, Morocco 109, 116, 123, 259, 261, 269, 272, 273 Tigris-Euphrates River 56, 86 till see glacial and periglacial environments; moraines tillage 182–3, 214 see also ploughing; soils TOMS (Total Ozone Mapping Spectrometer) 417–8, 419, 420 see also aeolian processes and landforms; aerosols Tortonian 11, 14, 15, 16, 19, 92, 94 tourism and recreation 70, 79, 83, 257, 258, 272, 273, 274, 276, 279, 280, 313, 314, 385, 404, 405, 406, 423, 425, 437, 438, 446, 448, 449, 457, 458, 459, 485, 523, 534, 554, 592, 593, 618, 621, 622, 632, 643 Tramontana wind 78, 531–2, 537 see also climate; storms and floods transgression, marine 23, 399, 401, 402, 403 see also sea-level change

transhumance 212, 563, 575, 637–8 see also pastoralism; grazing transpiration 204 see also evapotranspiration travertine 258, 297, 298, 305–6, 308, 315 see also tufa; speleothems Trieste, Italy 78, 288, 290, 301 Tripoli, Libya 288, 301, 401, 416, 422 Tripolitania, Libya 325, 327, 415 Troodos Mountains, Cyprus 70 Troy, Turkey 338, 485, 496, 504 tsunamis 20, 400–1, 437, 441, 446, 448–9, 455, 460, 470–1, 480, 482, 483–4, 493–510 quantification 493–4 major tsunami events in the Mediterranean 400–1, 494–505 tsunami generation 505 hazard assessment 505–8 risk mitigation 508–9 tufa 97, 261, 297, 298, 299, 302, 305–6, 308, 315, 321, 322 see also calcium carbonate; calcrete; travertine; speleothems Tunis 401, 416 Tunisia 10, 19, 33, 58, 75, 109, 116, 144, 150, 176, 213, 217, 232, 239, 248, 260, 274, 275, 276, 292, 315, 328, 330, 339, 360, 386, 393, 400, 401, 405, 415, 417, 418, 421, 423, 424, 425, 508, 515, 516, 530, 533, 534, 563, 571, 585, 586, 587, 589, 590, 591, 592, 594, 595, 596, 621, 631, 640, 642 turbidity and turbidity currents 20, 40, 346, 396, 460, 496, 586 see also turbidites turbidites and mega-turbidites 20, 400 Turkey 8, 10, 11, 18, 21, 39, 74, 75, 77, 80, 83, 84, 86, 102, 127, 159, 169, 171, 174, 176, 207, 212, 222, 232, 233, 234, 236, 255, 256, 257, 259, 260, 261, 262, 263, 266, 267, 268, 270, 271, 274, 275, 276, 278, 279, 287, 288, 293, 295, 298, 301, 305, 306, 308, 314, 319, 325, 331, 335, 354–5, 362, 364–5, 376, 385, 386, 390, 400, 405, 406, 418, 423, 435, 441, 454, 455, 469, 470, 471, 472, 477, 478, 480, 481, 486, 487, 499, 500, 504, 515, 533, 534, 535, 550, 571, 587, 588, 594, 595, 596, 615, 630, 631, 632, 640, 642, 643 Tuscany 13, 172, 182, 436, 442, 496, 497, 508, 516, 521, 621 Tyrrhenian Sea and coast 10, 13, 15, 36, 43, 102, 122, 215, 312, 388, 395,

Index 400, 448, 478, 496, 497, 506, 508, 515, 516, 521 Tyrrhenian basin 8, 9, 12, 13, 42, 95, 385, 387, 509 UN (United Nations) 565, 566 uplift (crustal) 7, 14, 15, 16, 17–8, 19, 21–3, 24, 36, 37, 77, 91, 103, 173, 189, 195, 197, 255, 261, 290, 292, 293, 302, 305, 325, 329, 333, 335, 389, 392, 397, 399, 401, 444, 446, 454, 456, 470, 473, 476, 477, 479, 480, 481, 482, 483, 498 see also tectonics and landscape development uranium-series (U/Th) dating 24, 103, 113, 303, 305, 306, 312, 315, 321, 323, 325, 328, 330, 331, 332, 333, 365, 368, 370, 376 urban environments and urbanization 79, 86, 240, 246, 260, 274, 276, 278, 279, 281, 313, 314, 345, 347, 404, 405, 424, 437, 449, 459, 485, 522, 523, 525, 529, 530, 535, 536, 555, 564, 586, 591, 596, 605, 609, 622, 625, 637, 643, 644 USA 18, 553 Vaiont dam disaster 179, 194–6 see also mass movements Valencia 249, 280, 423, 530, 641 Valencia basin and trough 7, 9, 10, 21, 385, 387, 401 valley fill sediments see river systems and environmental change valley floor environments 193, 235, 242, 302, 309, 321–3, 325, 327, 329, 331, 334–5, 337–8, 346–8, 641 see also deltas; floodplains; riparian habitats Van, Lake, Turkey 102, 123, 255, 257, 261, 263, 267–70, 281 van Andel, Tjeerd 337 varves 119, 257–60, 267–9, 272 see also laminated sediments vegetation and ecosystem dynamics 203–225, 567–70 vegetation history 89–128, 208, 213, 217, 222–4 see also pollen and pollen records Venice 75, 77, 83, 406, 531 Vera basin, Spain 24 Vesuvius 13, 435, 436, 437, 438, 441, 442, 444–6, 457–60 see also volcanoes and volcanism Vikos–Aoos National Park, Greece 640, 641 Vikos gorge, Greece 108, 212, 230, 235, 288, 296, 300, 302, 306, 323, 324, 640, 641

vines 183, 190, 214, 272, 572, 618 Voidomatis River, Greece 235, 296, 300, 323–4, 328, 330, 331–3, 377 volcanoes and volcanism 5, 7, 12–4, 17, 20, 22, 128, 139, 171–2, 319, 435–460 ash see tephra gas and aerosols 439–41 hazards 180, 269, 436–441 lake basins 89, 255, 258–63, 267 lava flows 438–9 provinces and geological setting 435–6, 441–57 pyroclastic density currents 438–9 risk management 457–460 tephra clouds and falls 438 tsunamis 20, 437, 441, 446, 448–9, 455, 460, 493, 497, 499–501, 505–6, 509 volcanism and: archaeological records 442–445 climate 79, 90, 128 coasts 398–400 mountain glaciation 355, 364 river behaviour 17, 335 waders 151–3, 637 see also bird fauna wadis 55, 99, 234, 325, 236, 327–9, 330, 405 see also river systems and environmental change warblers 147, 148–50, 152, 276, 618, 625 see also bird fauna Water Framework Directive (WFD) 250–51, 280, 321, 564, 579, 596 see also European Union water harvesting 239 see also agricultural terraces water resources 84, 86, 229, 235, 241, 247, 279, 280, 281, 299, 309, 313, 315, 345, 552, 563, 564, 565, 583–96 see also runoff competition and threats 590–1 geography 583–6 human impact on 586 management and sustainability 591–2 resource by country 586–8 resource by population 588–9 threat of climate change and temporal trends 592–6 water table 194, 211, 280, 295, 296, 297, 301, 304, 309, 312, 314, 346, 425, 442, 642 see also groundwater water quality 246, 250, 251, 274, 280, 281, 290, 296, 313, 314, 315, 437, 564, 586, 591, 592, 594 see also eutrophication wave action 77, 305, 385, 386, 394, 395, 396, 441, 450, 455, 460, 483,

663

484, 493–510, 532, 535 see also tsunamis; marine terraces; coastal environments wave cut platform see marine terraces weathering and rock breakdown 23, 102, 169–78, 197, 287, 292, 298, 300, 303, 305, 306, 307, 309, 310, 311, 321, 333, 366, 370, 416, 425 see also bedrock erosion and weathering; dissolution westerlies 36, 37, 102, 126, 602 see also climate Western Mediterranean Deep Water (WMDW) 34, 35, 40, 42, 43 see also Mediterranean Sea wetlands 110, 150, 151, 213, 232, 246, 255–267, 272–82, 345, 346, 402, 406, 571, 590, 603, 618, 621, 622, 627, 628, 629, 631, 640, 642, 643 see also lakes and wetlands wheat 160, 272, 639 White Mountains, Crete 216, 354 White Desert, Egypt, 305, 307 Wild boar (Sus scrofa) 146, 640 wildfires 541–56 controls on fire regime 548–9 extent of the problem 541–3 hazard and impact on ecosystems 550–53 hazard mitigation 553–5 historical context and data sources 546–8 magnitude and frequency 543–6 winds 78 see also climate, Bora; Etesian; Mistral; Sirocco; Tramontana wind storms 513, 515, 531–7 see also storms and floods Wolf (Canis lupus) 146, 154, 635, 640 Würm glaciation 141, 146, 362, 369–70, 373–5, 376 see also glacial and periglacial environments xerophytes 97, 576, 625–6 see also halophytes xerosols 176, 177, 178 see also soils Younger Fill 321–3, 335–336 see also river systems and environmental change Yugoslavia 11, 77, 85, 232, 272, 279, 298, 301, 314, 331, 355, 362, 365, 367, 386, 479, 541, 543, 638 Zagros Mountains 10, 12, 149, 255 Zewana, Wadi, Libya 325, 326–30 Zohary’s law 209, 222