Encyclopedia of Ocean Sciences, Second Edition, Volume 2: D-F

ENCYCLOPEDIA OF OCEAN SCIENCES SECOND EDITION Editor-in-chief JOHN H. STEELE Editors STEVE A. THORPE KARL K. TUREK...

Author: John H. Steele | Steve A. Thorpe | Karl K. Turekian

103 downloads 1760 Views 34MB Size Report

This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!
Report copyright / DMCA form

DOWNLOAD PDF

ENCYCLOPEDIA OF

OCEAN SCIENCES SECOND EDITION

Editor-in-chief

JOHN H. STEELE

Editors

STEVE A. THORPE KARL K. TUREKIAN

Boston • Heidelberg • London • New York • Oxford Paris • San Diego • San Francisco • Singapore • Sydney • Tokyo Academic Press is an imprint of Elsevier

(c) 2011 Elsevier Inc. All Rights Reserved.

ENCYCLOPEDIA OF

OCEAN SCIENCES SECOND EDITION


Subject Area Volumes from the Second Edition Climate & Oceans edited by Karl K. Turekian Elements of Physical Oceanography edited by Steve A. Thorpe Marine Biology edited by John H. Steele Marine Chemistry & Geochemistry edited by Karl K. Turekian Marine Ecological Processes edited by John H. Steele Marine Geology & Geophysics edited by Karl K. Turekian Marine Policy & Economics guest edited by Porter Hoagland, Marine Policy Center, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts Measurement Techniques, Sensors & Platforms edited by Steve A. Thorpe Ocean Currents edited by Steve A. Thorpe The Coastal Ocean edited by Karl K. Turekian The Upper Ocean edited by Steve A. Thorpe


ENCYCLOPEDIA OF

OCEAN SCIENCES SECOND EDITION Volume 2: D - F Editor-in-chief

JOHN H. STEELE

Editors

STEVE A. THORPE KARL K. TUREKIAN

Boston • Heidelberg • London • New York • Oxford Paris • San Diego • San Francisco • Singapore • Sydney • Tokyo Academic Press is an imprint of Elsevier


Academic Press is an imprint of Elsevier 32 Jamestown Road, London NW1 7BY, UK 30 Corporate Drive, Suite 400, Burlington, MA 01803, USA 525 B Street, Suite 1900, San Diego, CA 92101-4495, USA Copyright ^ 2009 Elsevier Ltd. All rights reserved

The following articles are US government works in the public domain and are not subject to copyright: Fish Predation and Mortality; International Organizations; Large Marine Ecosystems; Ocean Circulation: Meridional Overturning Circulation; Salt Marsh Vegetation; Satellite Passive-Microwave Measurements of Sea Ice; Satellite Oceanography, History and Introductory Concepts; Satellite Remote Sensing: Ocean Color; Science of Ocean Climate Models; Wind- and Buoyancy-Forced Upper Ocean. Fish Migration, Horizontal Crown Copyright 2001 Turbulence Sensors Canadian Crown Copyright 2001 No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means electronic, mechanical, photocopying, recording or otherwise without the prior written permission of the publisher

Permissions may be sought directly from Elsevier’s Science & Technology Rights Department in Oxford, UK: phone (+44)(0) 1865 843830; fax (+44)(0) 1865 853333; email: [email protected]. Alternatively you can submit your request online by visiting the Elsevier website at (http://elsevier.com/locate/permissions), and selecting Obtaining permissions to use Elsevier material

Notice No responsibility is assumed by the publisher for any injury and/or damage to persons or property as a matter of products liability, negligence or otherwise, or from any use or operation of any methods, products, instructions or ideas contained in the material herein, Because of rapid advances in the medical sciences, in particular, independent verification of diagnoses and drug dosages should be made

British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library

Library of Congress Control Number: 2009935908

ISBN: 978-0-12-375044-0

For information on all Elsevier publications visit our website at www.elsevierdirect.com

PRINTED AND BOUND IN ITALY 09 10 11 12 13 10 9 8 7 6 5 4 3 2 1


Editors

Editor-in-chief John H. Steele Marine Policy Center, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA

Editors Steve A. Thorpe National Oceanography Centre, University of Southampton Southampton, UK School of Ocean Sciences, University of Bangor, Menai Bridge, Anglesey, UK Karl K. Turekian Yale University, Department of Geology and Geophysics, New Haven, Connecticut, USA


v

Editorial Advisory Board John H. S. Blaxter Scottish Association for Marine Science Dunstaffnage Marine Laboratory Oban Argyll, UK Quentin Bone The Marine Biological Association of the United Kingdom Plymouth, UK Kenneth H. Brink Woods Hole Oceanographic Institution Woods Hole MA, USA Harry L. Bryden School of Ocean and Earth Science James Rennell Division University of Southampton Empress Dock Southampton, UK Robert Clark University of Newcastle upon Tyne Marine Sciences and Coastal Management Newcastle upon Tyne, UK J. Kirk Cochran State University of New York at Stony Brook Marine Sciences Research Center Stony Brook NY, USA Jeremy S. Collie Coastal Institute Graduate School of Oceanography University of Rhode Island South Ferry Road Narragansett RI, USA

Paul G. Falkowski Departments of Geological Sciences & Marine & Coastal Sciences Institute of Marine & Coastal Sciences School of Environmental & Biological Sciences Rutgers University New Brunswick NJ, USA Mike Fashamw Southampton Oceanography Centre University of Southampton Southampton UK John G. Field MArine REsearch (MA-RE) Institute University of Cape Town Rondebosch South Africa Michael Fogarty NOAA, National Marine Fisheries Service Woods Hole MA, USA Wilford D. Gardner Department of Oceanography Texas A&M University College Station TX, USA Ann Gargett Old Dominion University Center for Coastal Physical Oceanography Crittenton Hall Norfolk VA, USA

Peter J. Cook Australian Petroleum Cooperative Research Centre Canberra, Australia

vi

Robert A. Duce Departments of Oceanography and Atmospheric Sciences Texas A&M University College Station TX, USA

w

deceased


Editorial Advisory Board

Christopher Garrett University of Victoria Department of Physics Victoria British Columbia, Canada

Lindsay Lairdw Aberdeen University Zoology Department Aberdeen UK

W. John Gould Southampton Oceanography Centre University of Southampton Southampton UK

Peter S. Liss University of East Anglia School of Environmental Sciences Norwich, UK

John S. Grayw Institute of Marine Biology and Limnology University of Oslo Blindern Oslo, Norway

Ken Macdonald University of California Department of Geological Sciences Santa Barbara CA, USA

Gwyn Griffiths Southampton Oceanography Centre University of Southampton Southampton UK

Dennis McGillicuddy Woods Hole Oceanographic Institution Woods Hole MA, USA Alasdair McIntyre University of Aberdeen Department of Zoology Aberdeen UK

Stephen J. Hall World Fish Center Penang Malaysia Roger Harris Plymouth Marine Laboratory West Hoe Plymouth, UK Porter Hoagland Woods Hole Oceanographic Institution Woods Hole MA, USA George L. Hunt Jr. University of California, Irvine Department of Ecology and Evolutionary Biology Irvine CA, USA William J. Jenkins Woods Hole Oceanographic Institution Woods Hole MA, USA

w

deceased

W. Kendall Melville Scripps Institution of Oceanography UC San Diego La Jolla CA, USA John Milliman College of William and Mary School of Marine Sciences Gloucester Point VA, USA James N. Moum College of Oceanic and Atmospheric Sciences Oregon State University Corvallis OR, USA Michael M. Mullinw Scripps Institution of Oceanography Marine Life Research Group University of California San Diego La Jolla CA, USA


vii

viii

Editorial Advisory Board

Yoshiyuki Nozakiw University of Tokyo The Ocean Research Institute Nakano-ku Tokyo Japan

Ellen Thomas Yale University Department of Geology and Geophysics New Haven CT, USA

John Orcutt Scripps Institution of Oceanography Institute of Geophysics and Planetary Physics La Jolla CA, USA Richard F. Pittenger Woods Hole Oceanographic Institution Woods Hole MA, USA Gerold Siedler Universita¨t Kiel Institut fua¨r Meereskunde Kiel Germany

Peter L. Tyack Woods Hole Oceanographic Institution Woods Hole MA, USA Bruce A. Warren Woods Hole Oceanographic Institution Woods Hole MA, USA Wilford F. Weeks University of Alaska Fairbanks Department of Geology and Geophysics Fairbanks AK, USA

Robert C. Spindel University of Washington Applied Physics Laboratory Seattle WA, USA

Robert A. Weller Woods Hole Oceanographic Institution Woods Hole MA, USA

Colin P. Summerhayes Scientific Committee on Antarctic Research (SCAR) Scott Polar Institute Cambridge, UK

w

Stewart Turner Australian National University Research School of Earth Sciences Canberra Australia

James A. Yoder Woods Hole Oceanographic Institution Woods Hole MA, USA

deceased


Preface to Second Print Edition The first edition of the Encyclopedia of Ocean Sciences, published in print form in 2001, has proven to be a valuable asset for the marine science community – and more generally. The continuing rapid increase in electronic access to academic material led us initially to publish the second edition electronically. We have now added this print version of the second edition because of a demonstrated need for such a product. The encyclopedia can now be accessed in print or electronic format according to the preferences and needs of individuals and institutions. In this edition there are 54 new articles, 67 revisions of previous articles, and a completely revised and improved index. We are grateful to the members of the Editorial Advisory Board, nearly all of whom have stayed with us during the lengthy process of going electronic. The transition from Academic Press to Elsevier occurred between the two editions. We thank Dr. Debbie Tranter of Elsevier for her efforts to see this edition through its final stages.

Preface to First Edition In 1942, a monumental volume was published on The Oceans by H. U. Sverdrup, M. W. Johnson, and R. H. Fleming. It was comprehensive and covered the knowledge at that time of the scientific study of the oceans. This seminal book helped to initiate the tremendous burgeoning of marine research that occurred during the following decades. The Encyclopedia of Ocean Sciences aims to embody the great growth of knowledge in a major new reference work. There have been remarkable new approaches to the study of the oceans that blur the distinctions between the physical, chemical, biological, and geological disciplines. New theories and technologies have expanded our knowledge of ocean processes. For example, plate tectonics has revolutionized our view not only of the geology and geophysics of the seafloor but also of ocean chemistry and biology. Satellite remote sensing provides a global vision as well as detailed understanding of the close coupling of ocean physics and biology at local and regional scales. Exploration, fishing, warfare, and the impact of storms have driven the past study of the seas, but we now have a great public awareness of and concern with broader social and economic issues affecting the oceans. For this reason, we have invited articles explicitly on marine policy and environmental topics, as well as encouraged authors to address these aspects of their particular subjects. We believe the encyclopedia should be of use to those involved with policy and management as well as to students and researchers. Over 400 scientists have contributed to this description of what we now know about the oceans. They are distinguished researchers who have generously shared their knowledge of this ever-growing body of science. We are extremely grateful to all these authors, whose ability to write concisely on complex subjects has generated a perspective on our science that we, as editors, believe will enhance the appreciation of the oceans, their uses, and the research ahead. It has been a major challenge for the members of the Editorial Advisory Board to cover such a heterogeneous subject. Their knowledge of the diverse areas of research has guaranteed comprehensive coverage of the ocean sciences. The Board contributed significantly by suggesting topics, persuading authors to contribute, and reviewing drafts. Many of them wrote Overviews that give broad descriptions of major parts of the ocean sciences. Clearly, it was the dedicated involvement of the Editorial Advisory Board that made this venture successful. Such a massive enterprise as a multivolume encyclopedia would not be possible without the long-term commitment of the staff of the Major Reference Works team at Academic Press. In particular, we are very grateful for the consistent support of our Senior Developmental Editor, Colin McNeil, who has worked so well with us throughout the whole process. Also, we are very pleased that new technology permits enhanced search and retrieval through the Internet. We believe this will make the encyclopedia much more accessible to individual researchers and students.


ix

x

Preface to Second Print Edition

In Memoriam During the creation of the Encyclopedia of Ocean Sciences and also in several cases prior to the publication of the electronic Second Edition, several Associate Editors or designated Associate Editors died. We specifically acknowledge their role in making this work an effective publication. They are Mike Fasham, John S. Gray, Lindsay Laird, Michael Mullin and Yoshiyuki Nozaki. J. H. Steele, S. A. Thorpe, and K. K. Turekian Editors


Guide to Use of the Encyclopedia

Introductory Points In devising the vision and structure for the Encyclopedia, the Editors have striven to unite and interrelate all current knowledge that can be designated ‘‘Ocean Sciences’’. To aid users of the Encyclopedia, this new reference work offers intuitive searching and extensive cross-linking of content. These features are explained in more detail below.

Structure of the Encyclopedia The material in the Encyclopedia is arranged as a series of articles in alphabetical order. To help you realize the full potential of the material in the Encyclopedia we have provided three features to help you find the topic of your choice.

1. Contents Lists Your first point of reference will probably be the contents list. The contents list appearing in each volume will provide you with the page number of the article. Alternatively you may choose to browse through a volume using the alphabetical order of the articles as your guide. To assist you in identifying your location within the Encyclopedia a running headline indicates the current article.

2. Cross References All of the articles in the encyclopedia have heen extensively cross referenced. The cross references, which appear at the end of each article, have heen provided at three levels: i. To indicate if a topic is discussed in greater detail elsewhere.

ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.

ii. To draw the reader’s attention to parallel discussions in other articles. ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.


xi

xii

Guide to Use of the Encyclopedia

iii. To indicate material that broadens the discussion.

ACOUSTICS, ARCTIC See also: Acoustics in Marine Sediments. Acoustic Noise. Acoustics, Shallow Water. Arctic Ocean Circulation. Bioacoustics. Ice–ocean interaction. Nepheloid Layers. North Atlantic Oscillation (NAO). Ocean Circulation: Meridional Overturning Circulation. Platforms: Autonomous Underwater Vehicles. Satellite Passive-Microwave Measurements of Sea Ice. Sea Ice. Sea Ice: Overview. Seals. Seismic Structure. Tomography. Under-Ice Boundary Layer. Water Types and Water Masses.

3. Index The index will provide you with the volume and page number where the material is to be located, and the index entries differentiate between material that is a whole article, is part of an article or is data presented in a table or figure. On the opening page of the index detailed notes are provided.

4. Appendices In addition to the articles that form the main body of the encyclopedia, there are a number of appendices which provide bathymetric charts and lists of data used throughout the encyclopedia. The appendices are located in volume 6, before the index.

5. Contributors A full list of contributors appears at the beginning of volume 1.


Contributors Volume 1 E E Adams

N Caputi

Massachusetts Institute of Technology, Cambridge, MA, USA

Fisheries WA Research Division, North Beach, WA, Australia

T Akal

C A Carlson

NATO SACLANT Undersea Research Centre, La Spezia, Italy

University of California, Santa Barbara, CA, USA H Chamley

R Arimoto New Mexico State University, Carlsbad, NM, USA

Universite´ de Lille 1, Villeneuve d’Ascq, France R Chester

J L Bannister The Western Australian Museum, Perth, Western Australia

Liverpool University, Liverpool, Merseyside, UK V Christensen University of British Columbia, Vancouver, BC, Canada

E D Barton University of Wales, Bangor, UK

J W Dacey

N R Bates Bermuda Biological Station for Research, St George’s, Bermuda, USA

Woods Hole Oceanographic Institution, Woods Hole, MA, USA R A Duce

A Beckmann

Texas A&M University, College Station, TX, USA

Alfred-Wegener-Institut fu¨r Polar und Meeresforschung, Bremerhaven, Germany

H W Ducklow

P S Bell

The College of William and Mary, Gloucester Point, VA, USA

Proudman Oceanographic Laboratory, Liverpool, UK I Dyer G Birnbaum

Marblehead, MA, USA

Alfred-Wegener-Institut fu¨r Polar und Meeresforschung, Bremerhaven, Germany

D W Dyrssen Gothenburg University, Go¨teborg, Sweden

B O Blanton The University of North Carolina at Chapel Hill, Chapel Hill, NC, USA E A Boyle Massachusetts Institute of Technology, Cambridge, MA, USA

S M Evans Newcastle University, Newcastle, UK I Everson Anglia Ruskin University, Cambridge, UK

P Boyle

J W Farrington

University of Aberdeen, Aberdeen, UK

Woods Hole Oceanographic Institution, MA, USA

D M Bush

M Fieux

State University of West Georgia, Carrollton, GA, USA

Universite´ Pierre et Marie Curie, Paris, France

K Caldeira

R A Fine

Stanford University, Stanford, CA, USA

University of Miami, Miami, FL, USA


xiii

xiv

Contributors

K G Foote Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA L François University of Lie`ge, Lie`ge, Belgium M A M Friedrichs Old Dominion University, Norfolk, VA, USA T Gaston National Wildlife Research Centre, Quebec, Canada J Gemmrich University of Victoria, Victoria, BC, Canada Y Godde´ris University of Lie`ge, Lie`ge, Belgium D R Godschalk University of North Carolina, Chapel Hill, NC, USA A J Gooday Southampton Oceanography Centre, Southampton, UK A L Gordon Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA D A Hansell University of Miami, Miami FL, USA L W Harding Jr University of Maryland, College Park, MD, USA R Harris Plymouth Marine Laboratory, Plymouth, UK P J Herring Southampton Oceanography Centre, Southampton, UK B M Hickey University of Washington, Seattle, WA, USA M A Hixon Oregon State University, Corvallis, OR, USA E E Hofmann Old Dominion University, Norfolk, VA, USA S Honjo Woods Hole Oceanographic Institution, Woods Hole, MA, USA D J Howell Newcastle University, Newcastle, UK J M Huthnance CCMS Proudman Oceanographic Laboratory, Wirral, UK B Ja¨hne University of Heidelberg, Heidelberg, Germany F B Jensen SACLANT Undersea Research Centre, La Spezia, Italy A John Sir Alister Hardy Foundation for Ocean Science, Plymouth, UK

C D Jones University of Washington, Seattle, WA, USA P F Kingston Heriot-Watt University, Edinburgh, UK W Krauss Institut fu¨r Meereskunde an der Universita¨t Kiel, Kiel, Germany W A Kuperman Scripps Institution of Oceanography, University of California, San Diego, CA, USA D Lal Scripps Institute of Oceanography, University of California San Diego, La Jolla, CA, USA C S Law Plymouth Marine Laboratory, The Hoe, Plymouth, UK W J Lindberg University of Florida, Gainesville, FL, USA J R E Lutjeharms University of Cape Town, Rondebosch, South Africa P Malanotte-Rizzoli Massachusetts Institute of Technology, Cambridge, MA, USA W R Martin Woods Hole Oceanographic Institution, Woods Hole, MA, USA R P Matano Oregon State University, Corvallis, OR, USA J W McManus University of Miami, Miami, FL, USA G M McMurtry University of Hawaii at Manoa, Honolulu, HI, USA R Melville-Smith Fisheries WA Research Division, North Beach, WA, Australia P N Mikhalevsky Science Applications International Corporation, McLean, VA, USA W D Miller University of Maryland, College Park, MD, USA D Monahan University of New Hampshire, Durham, NH, USA J C Moore University of California at Santa Cruz, Santa Cruz, CA, USA A Morel Universite´ Pierre et Marie Curie, Villefranche-sur-Mer, France R Narayanaswamy The University of Manchester, Manchester, UK


Contributors W J Neal Grand Valley State University, Allendale, MI, USA D Pauly University of British Columbia, Vancouver, BC, Canada J W Penn Fisheries WA Research Division, North Beach, WA, Australia L C Peterson University of Miami, Miami, FL, USA S G Philander Princeton University, Princeton, NJ, USA N J Pilcher Universiti Malaysia Sarawak, Sarawak, Malaysia O H Pilkey Duke University, Durham, NC, USA

xv

D H Shull Western Washington University, Bellingham, WA, USA D K Steinberg College of William and Mary, Gloucester Pt, VA, USA L Stramma University of Kiel, Kiel, Germany R N Swift NASA Goddard Space Flight Center, Wallops Island, VA, USA T Takahashi Lamont Doherty Earth Observatory, Columbia University, Palisades, NY, USA P D Thorne Proudman Oceanographic Laboratory, Liverpool, UK

A R Piola Universidad de Buenos Aires, Buenos Aires, Argentina J M Prospero University of Miami, Miami, FL, USA S Rahmstorf Potsdam Institute for Climate Impact Research, Potsdam, Germany P C Reid SAHFOS, Plymouth, UK G Reverdin LEGOS, Toulouse Cedex, France S R Rintoul CSIRO Antarctic Climate and Ecosystems Cooperative Research Centre, Hobart, TAS, Australia J M Roberts Scottish Association for Marine Science, Oban, UK P A Rona Rutgers University, New Brunswick, NJ, USA T C Royer Old Dominion University, Norfolk, VA, USA B Rudels Finnish Institute of Marine Research, Helsinki, Finland

P L Tyack Woods Hole Oceanographic Institution, Woods Hole, USA T Tyrrell National Oceanography Centre, Southampton, UK F E Werner The University of North Carolina at Chapel Hill, Chapel Hill, NC, USA E A Widder Harbor Branch Oceanographic Institution, Fort Pierce, FL, USA D J Wildish Fisheries and Oceans Canada, St. Andrews, NB, Canada A J Williams, III Woods Hole Oceanographic Institution, Woods Hole, MA, USA D K Woolf Southampton Oceanography Centre, Southampton, UK

W Seaman University of Florida, Gainesville, FL, USA

C W Wright NASA Goddard Space Flight Center, Wallops Island, VA, USA

F Sevilla, III, University of Santo Tomas, Manila,The Philippines

J D Wright Rutgers University, Piscataway, NJ, USA

L V Shannon University of Cape Town, Cape Town, South Africa

J R Young The Natural History Museum, London, UK

G I Shapiro University of Plymouth, Plymouth, UK

H J Zemmelink University of Groningen, Haren, The Netherlands

A D Short University of Sydney, Sydney, Australia

W Zenk Universita¨t Kiel, Kiel, Germany


xvi

Contributors

Volume 2 G P Arnold Centre for Environment, Fisheries & Aquaculture Science, Suffolk, UK

K Dyer University of Plymouth, Plymouth, UK

K M Bailey Alaska Fisheries Science Center, Seattle, WA, USA

M Elliott Institute of Estuarine and Coastal Studies, University of Hull, Hull, UK

J G Baldauf Texas A&M University, College Station, TX, USA

D M Farmer Institute of Ocean Sciences, Sidney, BC, Canada

J Bascompte CSIC, Seville, Spain

A V Fedorov Yale University, New Haven, CT, USA

A Belgrano Institute of Marine Research, Lysekil, Sweden

M J Fogarty Northeast Fisheries Science Center, National Marine Fisheries Service, Woods Hole, MA, USA

O A Bergstad Institute of Marine Research, Flødevigen His, Norway J H S Blaxter Scottish Association for Marine Science, Argyll, UK

R Fonteyne Agricultural Research Centre, Ghent, Oostende, Belgium

Q Bone The Marine Biological Association of the United Kingdom, Plymouth, UK

D J Fornari Woods Hole Oceanographic Institution, Woods Hole, USA

I Boyd University of St. Andrews, St. Andrews, UK

A E Gargett Old Dominion University, Norfolk, VA, USA

K M Brander DTU Aqua, Charlottenlund, Denmark and International Council for the Exploration of the Sea (ICES), Copenhagen, Denmark

C H Gibson University of California, San Diego, La Jolla, CA, USA

J N Brown Yale University, New Haven, CT, USA T K Chereskin University of California San Diego, La Jolla, CA, USA J S Collie Danish Institute for Fisheries Research, Charlottenlund, Denmark and University of Rhode Island, Narragansett, RI, USA G Cresswell CSIRO Marine Research, Tasmania, Australia

J D M Gordon Scottish Association for Marine Science, Argyll, UK J F Grassle Rutgers University, New Brunswick, New Jersey, USA S J Hall Flinders University, Adelaide, SA, Australia N Hanson University of St. Andrews, St. Andrews, UK P J B Hart University of Leicester, Leicester, UK

J Davenport University College Cork, Cork, Ireland

K R Helfrich Woods Hole Oceanographic Institution, Woods Hole, MA, USA

R H Douglas City University, London, UK

D M Higgs University of Windsor, Windsor, ON, Canada

S Draxler Karl-Franzens-Universita¨t Graz, Graz, Austria

N G Hogg Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J T Duffy-Anderson Alaska Fisheries Science Center, Seattle, WA, USA J A Dunne Santa Fe Institute, Santa Fe, NM, USA and Pacific Ecoinformatics and Computational Ecology Lab, Berkely, CA, USA

E D Houde University of Maryland, Solomons, MD, USA V N de Jonge Department of Marine Biology, Groningen University, Haren, The Netherlands


Contributors K Katsaros Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA J M Klymak University of Victoria, Victoria, BC, Canada M Kucera Eberhard Karls Universita¨t Tu¨bingen, Tu¨bingen, Germany R S Lampitt University of Southampton, Southampton, UK J R N Lazier Bedford Institute of Oceanography, NS, Canada J R Ledwell Woods Hole Oceanographic Institution, Woods Hole, MA, USA P F J Lermusiaux Harvard University, Cambridge, MA, USA M E Lippitsch Karl-Franzens-Universita¨t Graz, Graz, Austria

xvii

T J Pitcher University of British Columbia, Vancouver, Canada A N Popper University of Maryland, College Park, MD, USA J F Price Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA R D Prien Southampton Oceanography Centre, Southampton, UK A-L Reysenbach Portland State University, Portland, OR, USA P L Richardson Woods Hole Oceanographic Institution, Woods Hole, MA, USA A R Robinson Harvard University, Cambridge, MA, USA M D J Sayer Dunstaffnage Marine Laboratory, Oban, Argyll, UK

B J McCay Rutgers University, New Brunswick, NJ, USA

R W Schmitt Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J D McCleave University of Maine, Orono, ME, USA

J Scott DERA Winfrith, Dorchester, Dorset, UK

D Minchin Marine Organism Investigations, Killaloe, Republic of Ireland

M P Sissenwine Northeast Fisheries Science Center, Woods Hole, MA, USA

C M Moore University of Essex, Colchester, UK K Moran University of Rhode Island, Narragansett, RI, USA G R Munro University of British Columbia, Vancouver, BC, Canada J D Nash Oregon State University, Corvallis, Oregon, OR, USA A C Naveira Garabato University of Southampton, Southampton, UK

T P Smith Northeast Fisheries Science Center, Woods Hole, MA, USA P V R Snelgrove Memorial University of Newfoundland, St John’s, NL, Canada M A Spall Woods Hole Oceanographic Institution, Woods Hole, MA, USA A Stigebrandt University of Gothenburg, Gothenburg, Sweden D A V Stow University of Southampton, Southampton, UK

J D Neilson Department of Fisheries and Oceans, New Brunswick, Canada

D J Suggett University of Essex, Colchester, UK

Y Nozakiw University of Tokyo, Tokyo, Japan

U R Sumaila University of British Columbia, Vancouver, BC, Canada

R I Perry Department of Fisheries and Oceans, British Columbia, Canada S G Philander Princeton University, Princeton, NJ, USA w

Deceased.

K S Tande Norwegian College of Fishery Science, Tromsø, Norway S A Thorpe National Oceanography Centre, Southampton, UK R S J Tol Economic and Social Research Institute, Dublin, Republic of Ireland


xviii

Contributors

K E Trenberth National Center for Atmospheric Research, Boulder, CO, USA J J Videler Groningen University, Haren, The Netherlands

R S Wells Chicago Zoological Society, Sarasota, FL, USA D C Wilson Institute for Fisheries Management and Coastal Community Development, Hirtshals, Denmark

Volume 3 S Ali Plymouth Marine Laboratory, Plymouth, UK

K H Coale Moss Landing Marine Laboratories, CA, USA

J T Andrews University of Colorado, Boulder, CO, USA

M F Coffins University of Texas at Austin, Austin, TX, USA

M A de Angelis Humboldt State University, Arcata, CA, USA

P J Corkeron James Cook University, Townsville, Australia

A J Arp Romberg Tiburon Center for Environment Studies, Tiburon, CA, USA

B C Coull University of South Carolina, Columbia, SC, USA

T Askew Harbor Branch Oceanographic Institute, Ft Pierce, FL, USA

R Cowen University of Miami, Miami, FL, USA

R D Ballard Institute for Exploration, Mystic, CT, USA

G Cresswell CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia

G Barnabe´ Universite´ de Montpellier II, France

D S Cronan Royal School of Mines, London, UK

R S K Barnes University of Cambridge, Cambridge, UK

J Csirke Food and Agriculture Organization of the United Nations, Rome, Italy

E D Barton University of Wales, Bangor, Menai Bridge, Anglesey, UK

G A Cutter Old Dominion University, Norfolk, VA, USA

D Bhattacharya University of Iowa, Iowa City, IA, USA

D J DeMaster North Carolina State University, Raleigh, NC, USA

F von Blanckenburg Universita¨t Bern, Bern, Switzerland

T D Dickey University of California, Santa Barbara, CA, USA

D R Bohnenstiehl North Carolina State University, Raleigh, NC, USA

D Diemand Coriolis, Shoreham, VT, USA

H L Bryden University of Southampton, Southampton, UK J Burger Rutgers University, Piscataway, NJ, USA S M Carbotte Lamont-Doherty Earth Observatory of Columbia University, Palisades, NY, USA G T Chandler University of South Carolina, Columbia, SC, USA M A Charette Woods Hole Oceanographic Institution, Woods Hole, MA, USA

C S M Doake British Antarctic Survey, Cambridge, UK C M Domingues CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia C J Donlon Space Applications Institute, Ispra, Italy F Doumenge Museé Oceánographique de Monaco, Monaco R A Dunn University of Hawaii at Manoa, Honolulu, HI, USA


Contributors R P Dziak Oregon State University/National Oceanic and Atmospheric Administration, Hatfield Marine Science Center, Newport, OR, USA O Eldholm University of Oslo, Oslo, Norway

xix

S K Hooker University of St. Andrews, St. Andrews, UK H Hotta Japan Marine Science & Technology Center, Japan G R Ierley University of California San Diego, La Jolla, CA, USA

A E Ellis Marine Laboratory, Aberdeen, Scotland, UK C R Engle University of Arkansas at Pine Bluff, Pine Bluff, AR, USA C C Eriksen University of Washington, Seattle, WA, USA V Ettwein University College London, London, UK S Farrow Carnegie Mellon University, Pittsburgh, PA, USA M Fieux Universite´ Pierre et Marie Curie, Paris Cedex, France N Forteath Inspection Head Wharf, TAS, Australia J D Gage Scottish Association for Marine Science, Oban, UK S M Garcia Food and Agriculture Organization of the United Nations, Rome, Italy

G Ito University of Hawaii at Manoa, Honolulu, HI, USA J Jacoby Woods Hole Oceanographic Institution, Woods Hole, MA, USA M J Kaiser Bangor University, Bangor, UK A E S Kemp University of Southampton, Southampton Oceanography Centre, Southampton, UK W M Kemp University of Maryland Center for Environmental Science, Cambridge, MD, USA V S Kennedy University of Maryland, Cambridge, MD, USA P F Kingston Heriot-Watt University, Edinburgh, UK G L Kooyman University of California San Diego, CA, USA

C Garrett University of Victoria, VIC, Canada

W Krijgsman University of Utrecht, Utrecht, The Netherlands

R N Gibson Scottish Association for Marine Science, Argyll, Scotland

J B Kristoffersen University of Bergen, Bergen, Norway

M Gochfeld Environmental and Community Medicine, Piscataway, NJ, USA

K Lambeck Australian National University, Canberra, ACT, Australia

H O Halvorson University of Massachusetts Boston, Boston, MA, USA

R S Lampitt University of Southampton, Southampton, UK

B U Haq Vendome Court, Bethesda, MD, USA

M Landry University of Hawaii at Manoa, Department of Oceanography, Honolulu, HI, USA

G R Harbison Woods Hole Oceanographic Institution, Woods Hole, MA, USA

C G Langereis University of Utrecht, Utrecht, The Netherlands A Lascaratos University of Athens, Athens, Greece

R M Haymon University of California, CA, USA

S Leibovich Cornell University, Ithaca, NY, USA

D L Hebert University of Rhode Island, RI, USA J E Heyning The Natural History Museum of Los Angeles County, Los Angeles, CA, USA P Hoagland Woods Hole Oceanographic Institution, Woods Hole, MA, USA

W G Leslie Harvard University, Cambridge, MA, USA C Llewellyn Plymouth Marine Laboratory, Plymouth, UK R A Lutz Rutgers University, New Brunswick, NJ, USA


xx

Contributors

K C Macdonald Department of Geological Sciences and Marine Sciences Institute, University of California, Santa Barbara, CA, USA F T Mackenzie University of Hawaii, Honolulu, HI, USA L P Madin Woods Hole Oceanographic Institution, Woods Hole, MA, USA M Maslin University College London, London, UK G A Maul Florida Institute of Technology, Melbourne, FL, USA M McNutt MBARI, Moss Landing, CA, USA M G McPhee McPhee Research Company, Naches, WA, USA A D Mclntyre University of Aberdeen, Aberdeen, UK J Mienert University of Tromsø, Tromsø, Norway G E Millward University of Plymouth, Plymouth, UK H Momma Japan Marine Science & Technology Center, Japan J H Morison University of Washington, Seattle, WA, USA A E Mulligan Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J E Petersen Oberlin College, Oberlin, OH, USA M Phillips Network of Aquaculture Centres in Asia-Pacific (NACA), Bangkok, Thailand B Qiu University of Hawaii at Manoa, Hawaii, USA F Quezada Biotechnology Center of Excellence Corporation, Waltham, MA, USA N N Rabalais Louisiana Universities Marine Consortium, Chauvin, LA, USA R D Ray NASA Goddard Space Flight Center, Greenbelt, MD, USA M R Reeve National Science Foundation, Arlington VA, USA R R Reeves Okapi Wildlife Associates, QC, Canada A Reyes-Prieto University of Iowa, Iowa City, IA, USA P B Rhines University of Washington,Seattle, WA, USA A R Robinson Harvard University, Cambridge, MA, USA H T Rossby University of Rhode Island, Kingston, RI, USA H M Rozwadowski Georgia Institute of Technology, Atlanta, Georgia, USA

W Munk University of California San Diego, La Jolla, CA, USA

A G V Salvanes University of Bergen, Bergen, Norway

E J Murphy British Antarctic Survey, Marine Life Sciences Division, Cambridge, UK

R Schlitzer Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany

P D Naidu National Institute of Oceanography, Dona Paula, India

M E Schumacher Woods Hole Oceanographic Institution, Woods Hole, MA, USA

N Niitsuma Shizuoka University, Shizuoka, Japan

M I Scranton State University of New York, Stony Brook, NY, USA

D B Olson University of Miami, Miami, FL, USA G-A Paffenho¨fer Skidaway Institute of Oceanography, Savannah, GA, USA C Paris University of Miami, Miami, FL, USA M R Perfit Department of Geological Sciences, University of Florida, Gainsville, FL, USA

K Sherman Narragansett Laboratory, Narragansett, RI, USA M D Spalding UNEP World Conservation Monitoring Centre and Cambridge Coastal Research Unit, Cambridge, UK J Sprintall University of California San Diego, La Jolla, CA, USA J H Steele Woods Hole Oceanographic Institution, MA, USA


Contributors C A Stein University of Illinois at Chicago, Chicago, IL, USA

S M Van Parijs Norwegian Polar Institute, Tromsø, Norway

C Stickley University College London, London, UK

L M Ver University of Hawaii, Honolulu, HI, USA

U R Sumaila University of British Columbia, Vancouver, BC, Canada

F J Vine University of East Anglia, Norwich, UK

S Takagawa Japan Marine Science & Technology Center, Japan

K L Von Damm University of New Hampshire, Durham, NH, USA

P K Taylor Southampton Oceanography Centre, Southampton, UK

R P Von Herzen Woods Hole Oceanographic Institution, Woods Hole, MA, USA

A Theocharis National Centre for Marine Research (NCMR), Hellinikon, Athens, Greece

xxi

D Wartzok Florida International University, Miami, FL, USA

P C Ticco Massachusetts Maritime Academy, Buzzards Bay, MA, USA R P Trask Woods Hole Oceanographic Institution, Woods Hole, MA, USA

W F Weeks Portland, OR, USA R A Weller Woods Hole Oceanographic Institution, Woods Hole, MA, USA

A W Trites University of British Columbia, British Columbia, Canada

J A Whitehead Woods Hole Oceanographic Institution, Woods Hole, MA, USA

A Turner University of Plymouth, Plymouth, UK

J C Wiltshire University of Hawaii, Manoa, Honolulu, HA, USA

P L Tyack Woods Hole Oceanographic Institution, Woods Hole, MA, USA

C Woodroffe University of Wollongong, Wollongong, NSW, Australia

G J C Underwood University of Essex, Colchester, UK

C Wunsch Massachusetts Institute of Technology, Cambridge, MA, USA

C L Van Dover The College of William and Mary, Williamsburg, VA, USA

H S Yoon University of Iowa, Iowa City, IA, USA

Volume 4 A Alldredge University of California, Santa Barbara, CA, USA D M Anderson Woods Hole Oceanographic Institution, Woods Hole, MA, USA O R Anderson Columbia University, Palisades, NY, USA

J M Bewers Bedford Institute of Oceanography, Dartmouth, NS, Canada N V Blough University of Maryland, College Park, MD, USA W Bonne

P G Baines CSIRO Atmospheric Research, Aspendale, VIC, Australia

Federal Public Service Health, Food Chain Safety and Environment, Brussels, Belgium

J M Baker Clock Cottage, Shrewsbury, UK

University of Cape Town, Cape Town, Republic of South Africa

J G Bellingham Monterey Bay Aquarium Research Institute, Moss Landing, CA, USA

R D Brodeur

G M Branch

Northwest Fisheries Science Center, Newport, OR, USA


xxii

Contributors

H Burchard Baltic Sea Research Institute Warnemu¨nde, Warnemu¨nde, Germany P H Burkill Plymouth Marine Laboratory, West Hoe, Plymouth, UK Francois Carlotti C.N.R.S./Universite´ Bordeaux 1, Arachon, France K L Casciotti Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J S Grayw University of Oslo, Oslo, Norway A G Grottoli University of Pennsylvania, Philadelphia, PA, USA N Gruber Institute of Biogeochemistry and Pollutant Dynamics, ETH Zurich, Switzerland K C Hamer University of Durham, Durham, UK D Hammond University of Southern California, Los Angeles, CA, USA

A Clarke British Antarctic Survey, Cambridge, UK

W W Hay Christian-Albrechts University, Kiel, Germany

M B Collins National Oceanography Centre, Southampton, UK

J W Heath Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA

J J Cullen Department of Oceanography, Halifax, NS, Canada D H Cushing Lowestoft, Suffolk, UK

D Hedgecock University of Southern California, Los Angeles, CA, USA C Hemleben Tu¨bingen University, Tu¨bingen, Germany

K L Denman University of Victoria, Victoria, BC, Canada S C Doney Woods Hole Oceanographic Institution, Woods Hole, MA, USA

T D Herbert Brown University, Providence, RI, USA I Hewson University of California Santa Cruz, Santa Cruz, CA, USA

J F Dower University of British Columbia, Vancouver, BC, Canada

Richard Hey University of Hawaii at Manoa, Honolulu, HI, USA

K Dysthe University of Bergen, Bergen, Norway

P Hoagland Woods Hole Oceanographic Institution, Woods Hole, MA, USA

H N Edmonds University of Texas at Austin, Port Aransas, TX, USA

N Hoepffner Institute for Environment and Sustainability, Ispra, Italy

L Føyn Institute of Marine Research, Bergen, Norway

M Hood Intergovernmental Oceanographic Commission, Paris, France

J Fuhrman University of Southern California, Los Angeles, CA, USA

M J Howarth Proudman Oceanographic Laboratory, Wirral, UK

C P Gallienne Plymouth Marine Laboratory, West Hoe, Plymouth, UK

M Huber Purdue University, West Lafayette, IN, USA

E Garel CIACOMAR, Algarve University, Faro, Portugal

J W Hurrell National Center for Atmospheric Research, Boulder, CO, USA

D M Glover Woods Hole Oceanographic Institution, Woods Hole, MA, USA S L Goodbred Jr State University of New York, Stony Brook, NY, USA J D M Gordon Scottish Association for Marine Science, Oban, Argyll, UK

D R Jackett CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia R A Jahnke Skidaway Institute of Oceanography, Savannah, GA, USA w

Deceased.


Contributors

xxiii

A Jarre University of Cape Town, Cape Town, South Africa

A F Michaels University of Southern California, Los Angeles, CA, USA

J Joseph La Jolla, CA, USA

J D Milliman College of William and Mary, Gloucester, VA, USA

D M Karl University of Hawaii at Manoa, Honolulu, HI, USA

C D Mobley Sequoia Scientific, Inc., WA, USA

K L Karsh Princeton University, Princeton, NJ, USA

M M Mullinw Scripps Institution of Oceanography, La Jolla, CA, USA

J Karstensen Universita¨t Kiel (IFM-GEOMAR), Kiel, Germany

P Mu¨ller University of Hawaii, Honolulu, HI, USA

R M Key Princeton University, Princeton, NJ, USA

L A Murray The Centre for Environment, Fisheries and Aquaculture Sciences, Lowestoft, UK

P D Killworth Southampton Oceanography Centre, Southampton, UK B Klinger Center for Ocean-Land-Atmosphere Studies (COLA), Calverton, MD, USA H E Krogstad NTNU, Trondheim, Norway I Laing Centre for Environment Fisheries and Aquaculture Science, Weymouth, UK

T Nagai Tokyo University of Marine Science and Technology, Tokyo, Japan K H Nisancioglu Bjerknes Centre for Climate Research, University of Bergen, Bergen, Norway Y Nozakiw University of Tokyo, Tokyo, Japan

G F Lane-Serff University of Manchester, Manchester, UK

K J Orians The University of British Columbia, Vancouver, BC, Canada

A Longhurst Place de I’Eglise, Cajarc, France

C A Paulson Oregon State University, Corvallis, OR, USA

R Lukas University of Hawaii at Manoa, Hawaii, USA

W G Pearcy Oregon State University, Corvallis, OR, USA

M Lynch University of California Santa Barbara, Santa Barbara, CA, USA

W S Pegau Oregon State University, Corvallis, OR, USA

M Macleod World Wildlife Fund, Washington, DC, USA E Maranõń University of Vigo, Vigo, Spain S Martin University of Washington, Seattle, WA, USA S M Masutani University of Hawaii at Manoa, Honolulu, HI, USA I N McCave University of Cambridge, Cambridge, UK T J McDougall CSIRO Marine and Atmospheric Research, Hobart, TAS, Australia C L Merrin The University of British Columbia, Vancouver, BC, Canada

T Platt Dalhousie University, NS, Canada J J Polovina National Marine Fisheries Service, Honolulu, HI, USA D Quadfasel Niels Bohr Institute, Copenhagen, Denmark J A Raven Biological Sciences, University of Dundee, Dundee, UK G E Ravizza Woods Hole Oceanographic Institution, Woods Hole, MA, USA A J Richardson University of Queensland, St. Lucia, QLD, Australia M Rubega University of Connecticut, Storrs, CT, USA w

Deceased.


xxiv

Contributors

K C Ruttenberg Woods Hole Oceanographic Institution, Woods Hole, MA, USA

K K Turekian Yale University, New Haven, CT, USA T Tyrrell University of Southampton, Southampton, UK

A G V Salvanes University of Bergen, Bergen, Norway

O Ulloa Universidad de Concepcioń, Concepcioń, Chile

S Sathyendranath Dalhousie University, NS, Canada

C M G Vivian The Centre for Environment, Fisheries and Aquaculture Sciences, Lowestoft, UK

R Schiebel Tu¨bingen University, Tu¨bingen, Germany F B Schwing NOAA Fisheries Service, Pacific Grove, CA, USA

J J Walsh University of South Florida, St. Petersburg, FL, USA

M P Seki National Marine Fisheries Service, Honolulu, HI, USA

R M Warwick Plymouth Marine Laboratory, Plymouth, UK

L J Shannon Marine and Coastal Management, Cape Town, South Africa

N C Wells Southampton Oceanography Centre, Southampton, UK

K Shepherd Institute of Ocean Sciences, Sidney, BC, Canada

J A Whitehead Woods Hole Oceanographic Institution, Woods Hole, MA, USA

D Siegel-Causey Harvard University, Cambridge MA, USA D M Sigman Princeton University, Princeton, NJ, USA A Soloviev Nova Southeastern University, FL, USA J H Steele Woods Hole Oceanographic Institution, MA, USA P K Takahashi University of Hawaii at Manoa, Honolulu, HI, USA L D Talley Scripps Institution of Oceanography, La Jolla, CA, USA E Thomas Yale University, New Haven, CT, USA J R Toggweiler NOAA, Princeton, NJ, USA

M Wilkinson Heriot-Watt University, Edinburgh, UK R G Williams University of Liverpool, Oceanography Laboratories, Liverpool, UK C A Wilson III Department of Oceanography and Coastal Sciences, and Coastal Fisheries Institute, CCEER Louisiana State University, Baton Rouge, LA, USA H Yamazaki Tokyo University of Marine Science and Technology, Tokyo, Japan B deYoung Memorial University, St. John’s, NL, Canada G Zibordi Institute for Environment and Sustainability, Ispra, Italy

Volume 5 D G Ainley H.T. Harvey Associates, San Jose CA, USA W Alpers University of Hamburg, Hamburg, Germany J R Apelw Global Ocean Associates, Silver Spring, MD, USA w

Deceased.

A B Baggeroer Massachusetts Institute of Technology, Cambridge, MA, USA L T Balance NOAA-NMFS, La Jolla, CA, USA R Batiza Ocean Sciences, National Science Foundation, VA, USA W H Berger Scripps Institution of Oceanography, La Jolla, CA, USA


Contributors J L Bodkin US Geological Survey, AK, USA

I Everson British Antarctic Survey Cambridge, UK

I L Boyd Natural Environment Research Council, Cambridge, UK

I Fer University of Bergen, Bergen, Norway

A C Brown University of Cape Town, Cape Town, Republic of South Africa

M Fieux Universite´-Pierre et Marie Curie, Paris, France

xxv

J Burger Rutgers University, Piscataway, NJ, USA

R A Flather Proudman Oceanographic Laboratory, Bidston Hill, Prenton, UK

C J Camphuysen Netherlands Institute for Sea Research, Texel, The Netherlands

G S Giese Woods Hole Oceanographic Institution, Woods Hole, MA, USA

D C Chapman Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J M Gregory Hadley Centre, Berkshire, UK

R E Cheney Laboratory for Satellite Altimetry, Silver Spring, Maryland, USA T Chopin University of New Brunswick, Saint John, NB, Canada J A Church Antarctic CRC and CSIRO Marine Research, TAS, Australia J K Cochran State University of New York, Stony Brook, NY, USA P Collar Southampton Oceanography Centre, Southampton, UK R J Cuthbert University of Otago, Dunedin, New Zealand L S Davis University of Otago, Dunedin, New Zealand K L Denman University of Victoria, Victoria BC, Canada R P Dinsmore Woods Hole Oceanographic Institution, Woods Hole, MA, USA G J Divoky University of Alaska, Fairbanks, AK, USA

S M Griffies NOAA/GFDL, Princeton, NJ, USA G Griffiths Southampton Oceanography Centre, Southampton, UK A Harding University of California, San Diego, CA, USA W S Holbrook University of Wyoming, Laramie, WY, USA G L Hunt, Jr University of Washington, Seattle, WA, USA and University of California, Irvine, CA, USA P Hutchinson North Atlantic Salmon Conservation Organization, Edinburgh, UK K B Katsaros Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA H L Kite-Powell Woods Hole Oceanographic Institution, Woods Hole, MA, USA M A Kominz Western Michigan University, Kalamazoo, MI, USA

L M Dorman University of California, San Diego, La Jolla, CA, USA

R G Kope Northwest Fisheries Science Center, Seattle, WA, USA

J F Dower University of British Columbia, Vancouver, BC, Canada

G S E Lagerloef Earth and Space Research, Seattle, WA, USA

J B Edson Woods Hole Oceanographic Institution, Woods Hole, MA, USA

L M Lairdw Aberdeen University, Aberdeen, UK

T I Eglinton Woods Hole Oceanographic Institution, Woods Hole, MA, USA

M Leppa¨ranta University of Helsinki, Helsinki, Finland w

Deceased.


xxvi

Contributors

E J Lindstrom NASA Science Mission Directorate, Washington, DC, USA

C T Roman University of Rhode Island, Narragansett, RI, USA

A K Liu NASA Goddard Space Flight Center, Greenbelt, MD, USA

M Sawhney University of New Brunswick, Saint John, NB, Canada

C R McClain NASA Goddard Space Flight Center, Greenbelt, MD, USA

G Shanmugam The University of Texas at Arlington, Arlington, TX, USA

D J McGillicuddy Jr Woods Hole Oceanographic Institution, Woods Hole, MA, USA

J Sharples Proudman Oceanographic Laboratory, Liverpool, UK

W K Melville Scripps Institution of Oceanography, La Jolla CA, USA

J H Simpson Bangor University, Bangor, UK R K Smedbol Dalhousie University, Halifax, NS, Canada

D Mills Atlantic Salmon Trust, UK

L B Spear H.T. Harvey Associates, San Jose, CA, USA

P J Minnett University of Miami, Miami, FL, USA W A Montevecchi Memorial University of Newfoundland, NL, Canada W S Moore University of South Carolina, Columbia, SC, USA S J Morreale Cornell University, Ithaca, NY, USA K W Nicholls British Antarctic Survey, Cambridge, UK T J O’Shea Midcontinent Ecological Science Center, Fort Collins, CO, USA T E Osterkamp University of Alaska, Alaska, AK, USA F V Paladino Indiana-Purdue University at Fort Wayne, Fort Wayne, IN, USA C L Parkinson NASA Goddard Space Flight Center, Greenbelt, MD, USA A Pearson Woods Hole Oceanographic Institution, Woods Hole, MA, USA J T Potemra SOEST/IPRC, University of Hawaii, Honolulu, HI, USA J A Powell Florida Marine Research Institute, St Petersburg, FL, USA T Qu SOEST/IPRC, University of Hawaii, Honolulu, HI, USA

R L Stephenson St. Andrews Biological Station, St. Andrews, NB, Canada J M Teal Woods Hole Oceanographic Institution, Rochester, MA, USA K K Turekian Yale University, New Haven, CT, USA P Wadhams University of Cambridge, Cambridge, UK W F Weeks Portland, OR, USA G Wefer Universita¨t Bremen, Bremen, Germany W S Wilson NOAA/NESDIS, Silver Spring, MD, USA M Windsor, North Atlantic Salmon Conservation Organization, Edinburgh, UK S Y Wu NASA Goddard Space Flight Center, Greenbelt, MD, USA L Yu Woods Hole Oceanographic Institution, Woods Hole, MA, USA


Contributors

xxvii

Volume 6 A V Babanin Swinburne University of Technology, Melbourne, VIC, Australia

S E Humphris Woods Hole Oceanographic Institution, Woods Hole, MA, USA

R T Barber Duke University Marine Laboratory, Beaufort, NC, USA

W J Jenkins University of Southampton, Southampton, UK

J Bartram World Health Organization, Geneva, Switzerland

D R B Kraemer The Johns Hopkins University, Baltimore, MD, USA

A Beckmann Alfred-Wegener-Institut fu¨r Polar- und Meeresforschung, Bremerhaven, Germany M C Benfield Louisiana State University, Baton Rouge, LA, USA P S Bogden Maine State Planning Office, Augusta, ME, USA J A T Bye The University of Melbourne, Melbourne, VIC, Australia M F Cronin NOAA Pacific Marine Environmental Laboratory, Seattle, WA, USA A R J David Bere Alston, Devon, UK W Deuser Woods Hole Oceanographic Institution, Woods Hole, MA, USA J Donat Old Dominion University, Norfolk, VA, USA C Dryden Old Dominion University, Norfolk, VA, USA A Dufour United States Environmental Protection Agency, OH, USA C A Edwards University of Connecticut, Groton, CT, USA W J Emery University of Colorado, Boulder, CO, USA E Fahrbach Alfred-Wegener-Institut fu¨r Polar- und Meeresforschung, Bremerhaven, Germany

S Krishnaswami Physical Research Laboratory, Ahmedabad, India E L Kunze University of Washington, Seattle, WA, USA T E L Langford University of Southampton, Southampton, UK J R Ledwell Woods Hole Oceanographic Institution, Woods Hole, MA, USA P L-F Liu Cornell University, Ithaca, NY, USA M M R van der Loeff Alfred-Wegener-Institut fu¨r Polar und Meereforschung Bremerhaven, Germany R Lueck University of Victoria, Victoria, BC, Canada J E Lupton Hatfield Marine Science Center, Newport, OR, USA L P Madin Woods Hole Oceanographic Institution, Woods Hole, MA, USA M E McCormick The Johns Hopkins University, Baltimore, MD, USA M G McPhee McPhee Research Company, Naches, WA, USA J H Middleton The University of New South Wales, Sydney, NSW, Australia P J Minnett University of Miami, Miami, FL, USA E C Monahan University of Connecticut at Avery Point, Groton, CT, USA

A M Gorlov Northeastern University, Boston, Massachusetts, USA

C Moore WET Labs Inc., Philomath, OR, USA

I Helmond CSIRO Marine Research, TAS, Australia

J H Morison University of Washington, Seattle, WA, USA

R A Holman Oregon State University, Corvallis, OR, USA

J N Moum Oregon State University, Corvallis, OR, USA


xxviii

Contributors

N S Oakey Bedford Institute of Oceanography, Dartmouth, NS, Canada D T Pugh University of Southampton, Southampton, UK D L Rudnick University of California, San Diego, CA, USA H Salas CEPIS/HEP/Pan American Health Organization, Lima, Peru L K Shay University of Miami, Miami, FL, USA W D Smyth Oregon State University, Corvallis, OR, USA J Sprintall University of California San Diego, La Jolla, CA, USA

L St. Laurrent University of Victoria, Victoria, BC, Canada W G Sunda National Ocean Service, NOAA, Beaufort, NC, USA M Tomczak Flinders University of South Australia, Adelaide, SA, Australia A J Watson University of East Anglia, Norwich, UK P H Wiebe Woods Hole Oceanographic Institution, Woods Hole, MA, USA P F Worcester University of California at San Diego, La Jolla, CA, USA


Contents Volume 1 Abrupt Climate Change

S Rahmstorf

1

Absorbance Spectroscopy for Chemical Sensors Abyssal Currents

R Narayanaswamy, F Sevilla, III

W Zenk

Accretionary Prisms

15

J C Moore

31

Acoustic Measurement of Near-Bed Sediment Transport Processes Acoustic Noise

Acoustic Scintillation Thermography Acoustics In Marine Sediments

K G Foote

62

P A Rona, C D Jones

71

T Akal

75

P N Mikhalevsky

Acoustics, Deep Ocean

92

W A Kuperman

101

F B Jensen

112

Acoustics, Shallow Water R Chester

Agulhas Current

120

J R E Lutjeharms

Aircraft Remote Sensing

128

L W Harding Jr, W D Miller, R N Swift, C W Wright

Air–Sea Gas Exchange

38 52

Acoustic Scattering by Marine Organisms

Aeolian Inputs

P D Thorne, P S Bell

I Dyer

Acoustics, Arctic

7

B Ja¨hne

138 147

Air–Sea Transfer: Dimethyl Sulfide, COS, CS2, NH4, Non-Methane Hydrocarbons, Organo-Halogens J W Dacey, H J Zemmelink

157

Air–Sea Transfer: N2O, NO, CH4, CO

163

Alcidae

C S Law

T Gaston

171

Antarctic Circumpolar Current Antarctic Fishes

S R Rintoul

I Everson

191

Anthropogenic Trace Elements in the Ocean Antifouling Materials

E A Boyle

211

W Seaman, W J Lindberg

234

R A Duce

Atmospheric Transport and Deposition of Particulate Material to the Oceans R Arimoto Authigenic Deposits

226

S G Philander

Atmospheric Input of Pollutants

Baleen Whales

203

B Rudels

Atlantic Ocean Equatorial Currents

Bacterioplankton

195

D J Howell, S M Evans

Arctic Ocean Circulation Artificial Reefs

178

G M McMurtry H W Ducklow

J L Bannister

238 J M Prospero, 248 258 269 276


xxix

xxx

Contents

Baltic Sea Circulation Bathymetry

W Krauss

288

D Monahan

297

Beaches, Physical Processes Affecting Benguela Current

Benthic Foraminifera

316 D J Wildish

328

A J Gooday

Benthic Organisms Overview

336

P F Kingston

348

P L Tyack

357

Biogeochemical Data Assimilation

E E Hofmann, M A M Friedrichs

Biological Pump and Particle Fluxes Bioluminescence

Bioturbation

S Honjo

376

A Morel

385

D H Shull

Black Sea Circulation

395

G I Shapiro

Bottom Water Formation

401

A L Gordon

415

Brazil and Falklands (Malvinas) Currents

A R Piola, R P Matano

Breaking Waves and Near-Surface Turbulence

J Gemmrich

D K Woolf

Calcium Carbonates

L C Peterson

E D Barton

Carbon Dioxide (CO2) Cycle

467

T Takahashi

Cenozoic Climate – Oxygen Isotope Evidence Cenozoic Oceans – Carbon Cycle Models

J D Wright L François, Y Godde´ris

R A Fine W R Martin

J W Farrington

Coastal Zone Management

514

539 551 563

F E Werner, B O Blanton

Coastal Topography, Human Impact on Coastal Trapped Waves

502

531

H Chamley

Coastal Circulation Models

495

524

Chemical Processes in Estuarine Sediments

Coccolithophores

E E Adams, K Caldeira

P Boyle

Chlorinated Hydrocarbons

477 487

Carbon Sequestration via Direct Injection into the Ocean

Clay Mineralogy

455

C A Carlson, N R Bates, D A Hansell, D K Steinberg

CFCs in the Ocean

431

445

B M Hickey, T C Royer

Canary and Portugal Currents

Cephalopods

422

439

California and Alaska Currents

Carbon Cycle

364 371

P J Herring, E A Widder

Bio-Optical Models

Bubbles

305

L V Shannon

Benthic Boundary Layer Effects

Bioacoustics

A D Short

D M Bush, O H Pilkey, W J Neal

J M Huthnance D R Godschalk

T Tyrrell, J R Young


572 581 591 599 606

Contents

Cold-Water Coral Reefs Conservative Elements

J M Roberts

615

D W Dyrssen

626

Continuous Plankton Recorders Copepods

A John, P C Reid

R Harris

Coral Reefs

630 640

Coral Reef and Other Tropical Fisheries Coral Reef Fishes

xxxi

V Christensen, D Pauly

M A Hixon

655

J W McManus

660

Corals and Human Disturbance Cosmogenic Isotopes

N J Pilcher

671

D Lal

678

Coupled Sea Ice–Ocean Models Crustacean Fisheries

651

A Beckmann, G Birnbaum

688

J W Penn, N Caputi, R Melville-Smith

699

CTD (Conductivity, Temperature, Depth) Profiler Current Systems in the Atlantic Ocean Current Systems in the Indian Ocean

A J Williams, III

L Stramma M Fieux, G Reverdin

Current Systems in the Southern Ocean

A L Gordon

Current Systems in the Mediterranean Sea

P Malanotte-Rizzoli

708 718 728 735 744

Volume 2 Data Assimilation in Models Deep Convection

A R Robinson, P F J Lermusiaux

J R N Lazier

Deep Submergence, Science of

13 D J Fornari

22

K Moran

37

Deep-Sea Drilling Methodology Deep-Sea Drilling Results

1

J G Baldauf

45

Deep-Sea Fauna

P V R Snelgrove, J F Grassle

55

Deep-Sea Fishes

J D M Gordon

67

Deep-Sea Ridges, Microbiology

A-L Reysenbach

73

Deep-Sea Sediment Drifts

D A V Stow

80

Demersal Species Fisheries

K Brander

90

Determination of Past Sea Surface Temperatures Differential Diffusion

A E Gargett

Dispersion from Hydrothermal Vents Diversity of Marine Species Dolphins and Porpoises

R W Schmitt, J R Ledwell

K R Helfrich

P V R Snelgrove R S Wells

Double-Diffusive Convection

98 114

Dispersion and Diffusion in the Deep Ocean

Drifters and Floats

M Kucera

R W Schmitt

P L Richardson


122 130 139 149 162 171

xxxii

Contents

Dynamics of Exploited Marine Fish Populations East Australian Current

M J Fogarty

G Cresswell

179 187

Economics of Sea Level Rise

R S J Tol

197

Ecosystem Effects of Fishing

S J Hall

201

Eels

J D McCleave

208

Effects of Climate Change on Marine Mammals Ekman Transport and Pumping

T K Chereskin, J F Price

El Ninõ Southern Oscillation (ENSO)

Electrical Properties of Sea Water

Energetics of Ocean Mixing

228

S G Philander

R D Prien

Elemental Distribution: Overview

Y Nozaki

255

A C Naveira Garabato

261

Eutrophication

271

J M Klymak, J D Nash

Estuarine Circulation

288

K Dyer

299

V N de Jonge, M Elliott

Evaporation and Humidity

Fiord Circulation

306

K Katsaros

Exotic Species, Introduction of Expendable Sensors

241 247

w

A V Fedorov, J N Brown

Estimates of Mixing

218 222

K E Trenberth

El Ninõ Southern Oscillation (ENSO) Models

Equatorial Waves

I Boyd, N Hanson

324

D Minchin

332

J Scott

345

A Stigebrandt

353

Fiordic Ecosystems

K S Tande

359

Fish Ecophysiology

J Davenport

367

Fish Feeding and Foraging Fish Larvae

P J B Hart

E D Houde

Fish Locomotion

381

J J Videler

Fish Migration, Horizontal Fish Migration, Vertical

Fish Reproduction

Fish Vision

392

G P Arnold

402

J D Neilson, R I Perry

Fish Predation and Mortality

Fish Schooling

374

411

K M Bailey, J T Duffy-Anderson

J H S Blaxter

425

T J Pitcher

432

R H Douglas

445

Fish: Demersal Fish (Life Histories, Behavior, Adaptations) Fish: General Review

O A Bergstad

Q Bone

458 467

Fish: Hearing, Lateral Lines (Mechanisms, Role in Behavior, Adaptations to Life Underwater) A N Popper, D M Higgs w

417

Deceased.


476

Contents

Fisheries and Climate

K M Brander

Fisheries Economics

483

U R Sumaila, G R Munro

Fisheries Overview

491

M J Fogarty, J S Collie

Fisheries: Multispecies Dynamics Fishery Management

499

J S Collie

505

T P Smith, M P Sissenwine

Fishery Management, Human Dimension

513

D C Wilson, B J McCay

Fishery Manipulation through Stock Enhancement or Restoration Fishing Methods and Fishing Fleets Floc Layers

xxxiii

M D J Sayer

R Fonteyne

522 528 535

R S Lampitt

548

Florida Current, Gulf Stream, and Labrador Current Flow through Deep Ocean Passages Flows in Straits and Channels

P L Richardson

N G Hogg

554 564

D M Farmer

572

Fluid Dynamics, Introduction, and Laboratory Experiments

S A Thorpe

578

Fluorometry for Biological Sensing

D J Suggett, C M Moore

581

Fluorometry for Chemical Sensing

S Draxler, M E Lippitsch

589

Food Webs

A Belgrano, J A Dunne, J Bascompte

Forward Problem in Numerical Models Fossil Turbulence

596

M A Spall

604

C H Gibson

612

Volume 3 Gas Exchange in Estuaries

M I Scranton, M A de Angelis

Gelatinous Zooplankton

L P Madin, G R Harbison

General Circulation Models

Geomorphology

20

C G Langereis, W Krijgsman

C Woodroffe

Geophysical Heat Flow

C A Stein, R P Von Herzen

40 K Lambeck

C C Eriksen A D Mclntyre

Grabs for Shelf Benthic Sampling

67

P F Kingston

70

M McNutt

80

Groundwater Flow to the Coastal Ocean Habitat Modification

49 59

Global Marine Pollution

Gravity

25 33

Glacial Crustal Rebound, Sea Levels, and Shorelines Gliders

9

G R Ierley

Geomagnetic Polarity Timescale

1

A E Mulligan, M A Charette

M J Kaiser

Heat and Momentum Fluxes at the Sea Surface Heat Transport and Climate History of Ocean Sciences

88 99

P K Taylor

H L Bryden H M Rozwadowski


105 114 121

xxxiv

Contents

Holocene Climate Variability Hydrothermal Vent Biota

M Maslin, C Stickley, V Ettwein R A Lutz

125 133

Hydrothermal Vent Deposits

R M Haymon

144

Hydrothermal Vent Ecology

C L Van Dover

151

Hydrothermal Vent Fauna, Physiology of

A J Arp

159

Hydrothermal Vent Fluids, Chemistry of

K L Von Damm

164

Hypoxia

N N Rabalais

172

Icebergs

D Diemand

181

Ice-Induced Gouging of the Seafloor Ice–Ocean Interaction

W F Weeks

J H Morison, M G McPhee

191 198

Ice Shelf Stability

C S M Doake

209

Igneous Provinces

M F Coffins, O Eldholm

218

Indian Ocean Equatorial Currents Indonesian Throughflow

M Fieux

J Sprintall

237

Inherent Optical Properties and Irradiance Internal Tidal Mixing Internal Tides

T D Dickey

W Munk

258

C Garrett

266

International Organizations Intertidal Fishes

M R Reeve

274

R N Gibson

Intra-Americas Sea

244 254

R D Ray

Internal Waves

Intrusions

226

280

G A Maul

286

D L Hebert

295

Inverse Modeling of Tracers and Nutrients

R Schlitzer

300

Inverse Models

C Wunsch

312

IR Radiometers

C J Donlon

319

K H Coale

331

Iron Fertilization Island Wakes Krill

E D Barton

343

E J Murphy

349

Kuroshio and Oyashio Currents

B Qiu

Laboratory Studies of Turbulent Mixing Lagoons

358 J A Whitehead

R S K Barnes

Lagrangian Biological Models Land–Sea Global Transfers

377 D B Olson, C Paris, R Cowen F T Mackenzie, L M Ver

Langmuir Circulation and Instability Large Marine Ecosystems

S Leibovich

K Sherman

Laridae, Sternidae, and Rynchopidae Law of the Sea

371

389 394 404 413

J Burger, M Gochfeld

P Hoagland, J Jacoby, M E Schumacher (c) 2011 Elsevier Inc. All Rights Reserved.

420 432

Contents

Leeuwin Current

G Cresswell, C M Domingues

Long-Term Tracer Changes Macrobenthos Magnetics

444

F von Blanckenburg

455

J D Gage

467

F J Vine

478

Manganese Nodules Mangroves

xxxv

D S Cronan

488

M D Spalding

496

Manned Submersibles, Deep Water

H Hotta, H Momma, S Takagawa

Manned Submersibles, Shallow Water

T Askew

505 513

Mariculture Diseases and Health

A E Ellis

519

Mariculture of Aquarium Fishes

N Forteath

524

Mariculture of Mediterranean Species Mariculture Overview

G Barnabe´, F Doumenge

M Phillips

537

Mariculture, Economic and Social Impacts Marine Algal Genomics and Evolution Marine Biotechnology

532

C R Engle

545

A Reyes-Prieto, H S Yoon, D Bhattacharya

H O Halvorson, F Quezada

552 560

Marine Chemical and Medicine Resources

S Ali, C Llewellyn

567

Marine Fishery Resources, Global State of

J Csirke, S M Garcia

576

Marine Mammal Diving Physiology

G L Kooyman

Marine Mammal Evolution and Taxonomy

J E Heyning

Marine Mammal Migrations and Movement Patterns Marine Mammal Overview

582

P J Corkeron, S M Van Parijs

P L Tyack

Marine Mammal Trophic Levels and Interactions Marine Mammals and Ocean Noise

A W Trites

Marine Policy Overview Marine Protected Areas Marine Silica Cycle

635 643

654 G-A Paffenho¨fer

656

P Hoagland, P C Ticco

664

P Hoagland, U R Sumaila, S Farrow D J DeMaster

R S Lampitt

Maritime Archaeology

622

651

J H Steele

Marine Plankton Communities

Mediterranean Sea Circulation

672 678 686

R D Ballard

Meddies and Sub-Surface Eddies

Meiobenthos

S K Hooker

A E S Kemp

Marine Mesocosms

615

628

R R Reeves

Marine Mammals: Sperm Whales and Beaked Whales

Marine Snow

P L Tyack

D Wartzok

Marine Mammals, History of Exploitation

596 605

Marine Mammal Social Organization and Communication

Marine Mats

589

H T Rossby A R Robinson, W G Leslie, A Theocharis, A Lascaratos

B C Coull, G T Chandler (c) 2011 Elsevier Inc. All Rights Reserved.

695 702 710 726

xxxvi

Contents

Mesocosms: Enclosed Experimental Ecosystems in Ocean Science Mesopelagic Fishes

J E Petersen, W M Kemp

A G V Salvanes, J B Kristoffersen

Mesoscale Eddies

748

P B Rhines

Metal Pollution

755

G E Millward, A Turner

Metalloids and Oxyanions

732

768

G A Cutter

776

Methane Hydrates and Climatic Effects

B U Haq

784

Methane Hydrate and Submarine Slides

J Mienert

790

Microbial Loops

M Landry

Microphytobenthos

799

G J C Underwood

807

Mid-Ocean Ridge Geochemistry and Petrology Mid-Ocean Ridge Seismic Structure Mid-Ocean Ridge Seismicity

M R Perfit

815

S M Carbotte

826

D R Bohnenstiehl, R P Dziak

Mid-Ocean Ridge Tectonics, Volcanism, and Geomorphology

837 K C Macdonald

Mid-Ocean Ridges: Mantle Convection and Formation of the Lithosphere Millennial-Scale Climate Variability

J T Andrews

Mineral Extraction, Authigenic Minerals Molluskan Fisheries Monsoons, History of Moorings

G Ito, R A Dunn

852 867 881

J C Wiltshire

890

V S Kennedy

899

N Niitsuma, P D Naidu

910

R P Trask, R A Weller

919

Volume 4 Nekton

W G Pearcy, R D Brodeur

Nepheloid Layers

1

I N McCave

Network Analysis of Food Webs

8 J H Steele

Neutral Surfaces and the Equation of State Nitrogen Cycle

19 T J McDougall, D R Jackett

D M Karl, A F Michaels

Nitrogen Isotopes in the Ocean Noble Gases and the Cryosphere Non-Rotating Gravity Currents North Atlantic Oscillation (NAO) North Sea Circulation

25 32

D M Sigman, K L Karsh, K L Casciotti

40

M Hood

55

P G Baines

59

J W Hurrell

65

M J Howarth

73

Nuclear Fuel Reprocessing and Related Discharges

H N Edmonds

82

Ocean Biogeochemistry and Ecology, Modeling of

N Gruber, S C Doney

89

Ocean Carbon System, Modeling of Ocean Circulation

S C Doney, D M Glover

N C Wells

105 115

Ocean Circulation: Meridional Overturning Circulation

J R Toggweiler


126

Contents

Ocean Gyre Ecosystems

M P Seki, J J Polovina

Ocean Margin Sediments Ocean Ranching

132

S L Goodbred Jr

138

A G V Salvanes

146

R G Williams

156

Ocean Subduction

Ocean Thermal Energy Conversion (OTEC) Ocean Zoning

S M Masutani, P K Takahashi

M Macleod, M Lynch, P Hoagland

Offshore Sand and Gravel Mining Oil Pollution

Okhotsk Sea Circulation

E Garel, W Bonne, M B Collins

200

H Yamazaki, H Burchard, K Denman, T Nagai

Open Ocean Convection

A Soloviev, B Klinger

Open Ocean Fisheries for Deep-Water Species

Optical Particle Characterization

P H Burkill, C P Gallienne

265 272 274

R Lukas

287

E Thomas

295 W W Hay

Paleoceanography: Orbitally Tuned Timescales Paleoceanography: the Greenhouse World Particle Aggregation Dynamics Past Climate from Corals

T D Herbert

M Huber, E Thomas

A Alldredge

A G Grottoli

K L Denman, J F Dower

Pelagic Biogeography

A Longhurst

D H Cushing

Pelecaniformes

Peru–Chile Current System

C A Paulson, W S Pegau

J Karstensen, O Ulloa

319 330 338 348 356

379 385 393

K C Ruttenberg

Photochemical Processes

311

370

M Rubega

Phosphorus Cycle

303

364

D Siegel-Causey

Penetrating Shortwave Radiation

252 261

I Laing

Paleoceanography, Climate Models in

Phytobenthos

R A Jahnke

K K Turekian

Pacific Ocean Equatorial Currents

Phalaropes

243

G F Lane-Serff

Oysters – Shellfish Farming

Pelagic Fishes

234

K K Turekian

Oxygen Isotopes in the Ocean

Paleoceanography

226

J Joseph

Organic Carbon Cycling in Continental Margin Environments

Overflows and Cascades

208 218

J D M Gordon

Open Ocean Fisheries for Large Pelagic Species

Origin of the Oceans

182 191

L D Talley

One-Dimensional Models

167 174

J M Baker

Patch Dynamics

xxxvii

N V Blough

M Wilkinson

401 414 425


xxxviii

Contents

Phytoplankton Blooms

D M Anderson

Phytoplankton Size Structure Plankton

M M Mullin

Plankton and Climate Plankton Viruses

432

E Maranõń

445

w

453

A J Richardson

455

J Fuhrman, I Hewson

465

Platforms: Autonomous Underwater Vehicles Platforms: Benthic Flux Landers

J G Bellingham

R A Jahnke

485

Platinum Group Elements and their Isotopes in the Ocean

G E Ravizza

Plio-Pleistocene Glacial Cycles and Milankovitch Variability Polar Ecosystems

K H Nisancioglu

A Clarke

Pollution, Solids

494 504 514

C M G Vivian, L A Murray

Pollution: Approaches to Pollution Control Pollution: Effects on Marine Communities Polynyas

473

519

J S Grayw, J M Bewers R M Warwick

526 533

S Martin

540

Population Dynamics Models

Francois Carlotti

Population Genetics of Marine Organisms Pore Water Chemistry

546

D Hedgecock

556

D Hammond

Primary Production Distribution

563

S Sathyendranath, T Platt

572

Primary Production Methods

J J Cullen

578

Primary Production Processes

J A Raven

585

Procellariiformes

K C Hamer

590

Propagating Rifts and Microplates

Richard Hey

597

Protozoa, Planktonic Foraminifera

R Schiebel, C Hemleben

606

Protozoa, Radiolarians

O R Anderson

Radiative Transfer in the Ocean Radioactive Wastes Radiocarbon

C D Mobley

619

L Føyn

629

R M Key

637

Rare Earth Elements and their Isotopes in the Ocean Red Sea Circulation Redfield Ratio Refractory Metals

Y Nozaki

w

D Quadfasel

677

K J Orians, C L Merrin

Regime Shifts, Physical Forcing Regime Shifts: Methods of Analysis

653 666

T Tyrrell

Regime Shifts, Ecological Aspects

w

613

L J Shannon, A Jarre, F B Schwing F B Schwing B deYoung, A Jarre

Deceased.


687 699 709 717

Contents

Regional and Shelf Sea Models

J J Walsh

Remote Sensing of Coastal Waters

Rigs and offshore Structures River Inputs

722

N Hoepffner, G Zibordi

Remotely Operated Vehicles (ROVs)

xxxix

732

K Shepherd

742

C A Wilson III, J W Heath

748

J D Milliman

754

Rocky Shores

G M Branch

762

Rogue Waves

K Dysthe, H E Krogstad, P Mu¨ller

770

Rossby Waves

P D Killworth

Rotating Gravity Currents

781

J A Whitehead

790

Volume 5 Salmon Fisheries, Atlantic

P Hutchinson, M Windsor

Salmon Fisheries, Pacific Salmonid Farming Salmonids

1

R G Kope

L M Laird

12

w

23

D Mills

29

Salt Marsh Vegetation

C T Roman

Salt Marshes and Mud Flats Sandy Beaches, Biology of Satellite Altimetry

39

J M Teal

43

A C Brown

49

R E Cheney

58

Satellite Oceanography, History, and Introductory Concepts J R Apel w

W S Wilson, E J Lindstrom, 65

Satellite Passive-Microwave Measurements of Sea Ice

C L Parkinson

80

Satellite Remote Sensing of Sea Surface Temperatures

P J Minnett

91

Satellite Remote Sensing SAR

A K Liu, S Y Wu

Satellite Remote Sensing: Ocean Color

C R McClain

Satellite Remote Sensing: Salinity Measurements Science of Ocean Climate Models Sea Ice

103

G S E Lagerloef

S M Griffies

P Wadhams

114 127 133 141

Sea Ice Dynamics

M Leppa¨ranta

159

Sea Ice: Overview

W F Weeks

170

Sea Level Change

J A Church, J M Gregory

179

Sea Level Variations Over Geologic Time Sea Otters

w

M A Kominz

J L Bodkin

185 194

Deceased.


xl

Contents

Sea Surface Exchanges of Momentum, Heat, and Fresh Water Determined by Satellite Remote Sensing L Yu

202

Sea Turtles

212

F V Paladino, S J Morreale

Seabird Conservation

J Burger

Seabird Foraging Ecology Seabird Migration

220

L T Balance, D G Ainley, G L Hunt Jr

L B Spear

227 236

Seabird Population Dynamics

G L Hunt Jr

247

Seabird Reproductive Ecology

L S Davis, R J Cuthbert

251

Seabird Responses to Climate Change Seabirds and Fisheries Interactions

David G Ainley, G J Divoky C J Camphuysen

Seabirds as Indicators of Ocean Pollution Seabirds: An Overview Seals

265

W A Montevecchi

274

G L Hunt, Jr

279

I L Boyd

285

Seamounts and Off-Ridge Volcanism Seas of Southeast Asia

R Batiza

292

J T Potemra, T Qu

Seaweeds and their Mariculture Sediment Chronologies

305

T Chopin, M Sawhney

317

J K Cochran

327

Sedimentary Record, Reconstruction of Productivity from the Seiches

Seismic Structure

I Fer, W S Holbrook

L M Dorman

367

K B Katsaros

375

Sensors for Micrometeorological and Flux Measurements Shelf Sea and Shelf Slope Fronts

J B Edson

J Sharples, J H Simpson

H L Kite-Powell

Single Point Current Meters

T I Eglinton, A Pearson

P Collar, G Griffiths

436

Slides, Slumps, Debris Flows, and Turbidity Currents Small Pelagic Species Fisheries Small-Scale Patchiness, Models of

419 428

T J O’Shea, J A Powell G Shanmugam

R L Stephenson, R K Smedbol D J McGillicuddy Jr

Small-Scale Physical Processes and Plankton Biology

J F Dower, K L Denman

M Fieux

447 468 474 488 494

A B Baggeroer

Southern Ocean Fisheries

391

409

Single Compound Radiocarbon Measurements

Sonar Systems

382

401

R P Dinsmore

Somali Current

351 361

Sensors for Mean Meteorology

Shipping and Ports

333 344

A Harding

Seismology Sensors

Sirenians

G Wefer, W H Berger

D C Chapman, G S Giese

Seismic Reflection Methods for Study of the Water Column

Ships

257

504

I Everson


513

Contents

Sphenisciformes

L S Davis

520

Stable Carbon Isotope Variations in the Ocean Storm Surges

K K Turekian

529

R A Flather

530

Sub Ice-Shelf Circulation and Processes Submarine Groundwater Discharge Sub-Sea Permafrost Surface Films

xli

K W Nicholls

541

W S Moore

551

T E Osterkamp

559

W Alpers

570

Surface Gravity and Capillary Waves

W K Melville

573

Volume 6 Temporal Variability of Particle Flux Thermal Discharges and Pollution

W Deuser

1

T E L Langford

10

Three-Dimensional (3D) Turbulence Tidal Energy Tides

W D Smyth, J N Moum

A M Gorlov

26

D T Pugh

Tomography

32

P F Worcester

Topographic Eddies Towed Vehicles

40

J H Middleton

57

I Helmond

Trace Element Nutrients

65

W G Sunda

Tracer Release Experiments

75

A J Watson, J R Ledwell

Tracers of Ocean Productivity

Transmissometry and Nephelometry Tritium–Helium Dating

87

W J Jenkins

93

Transition Metals and Heavy Metal Speciation

Tsunami

18

J Donat, C Dryden

100

C Moore

109

W J Jenkins

119

P L-F Liu

127

Turbulence in the Benthic Boundary Layer Turbulence Sensors

R Lueck, L St. Laurrent, J N Moum

N S Oakey

Under-Ice Boundary Layer

148

M G McPhee, J H Morison

Upper Ocean Heat and Freshwater Budgets Upper Ocean Mean Horizontal Structure Upper Ocean Mixing Processes

155

P J Minnett

163

M Tomczak

175

J N Moum, W D Smyth

185

Upper Ocean Structure: Responses to Strong Atmospheric Forcing Events Upper Ocean Time and Space Variability Upper Ocean Vertical Structure Upwelling Ecosystems

141

L K Shay

192

D L Rudnick

211

J Sprintall, M F Cronin

217

R T Barber

Uranium-Thorium Decay Series in the Oceans: Overview

225 M M R van der Loeff


233

xlii

Contents

Uranium-Thorium Series Isotopes in Ocean Profiles Vehicles for Deep Sea Exploration

S E Humphris

Viral and Bacterial Contamination of Beaches Volcanic Helium

J Bartram, H Salas, A Dufour

285

Water Types and Water Masses

W J Emery

291

M E McCormick, D R B Kraemer

Waves on Beaches

267 277

E L Kunze

Wave Generation by Wind

244 255

J E Lupton

Vortical Modes

Wave Energy

S Krishnaswami

300

J A T Bye, A V Babanin

304

R A Holman

310

Weddell Sea Circulation

E Fahrbach, A Beckmann

318

Wet Chemical Analyzers

A R J David

326

Whitecaps and Foam

E C Monahan

Wind- and Buoyancy-Forced Upper Ocean Wind Driven Circulation

331 M F Cronin, J Sprintall

P S Bogden, C A Edwards

Zooplankton Sampling with Nets and Trawls

337 346

P H Wiebe, M C Benfield

355

Appendix 1. SI Units and Some Equivalences

373

Appendix 2. Useful Values

376

Appendix 3. Periodic Table of the Elements

377

Appendix 4. The Geologic Time Scale

378

Appendix 5. Properties of Seawater

379

Appendix 6. The Beaufort Wind Scale and Seastate

384

Appendix 7. Estimated Mean Oceanic Concentrations of the Elements

386

Appendix 8. Abbreviations

389

Appendix 9. Taxonomic Outline Of Marine Organisms

L P Madin

401

Appendix 10. Bathymetric Charts of the Oceans

412

Index

421


DATA ASSIMILATION IN MODELS A. R. Robinson and P. F. J. Lermusiaux, Harvard University, Cambridge, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 623–634, & 2001, Elsevier Ltd.

Introduction Data assimilation is a novel, versatile methodology for estimating oceanic variables. The estimation of a quantity of interest via data assimilation involves the combination of observational data with the underlying dynamical principles governing the system under observation. The melding of data and dynamics is a powerful methodology which makes possible efficient, accurate, and realistic estimations otherwise not feasible. It is providing rapid advances in important aspects of both basic ocean science and applied marine technology and operations. The following sections introduce concepts, describe purposes, present applications to regional dynamics and forecasting, overview formalism and methods, and provide a selected range of examples.

Field and Parameter Estimation Ocean science, and marine technology and operations, require a knowledge of the distribution and evolution in space and time of the properties of the sea. The functions of space and time, or state variables, which characterize the state of the sea under observation are classically designated as fields. The determination of state variables poses problems of state estimation or field estimation in three or four dimensions. The fundamental problem of ocean science may be simply stated as follows: given the state of the ocean at one time, what is the state of the ocean at a later time? It is the dynamics, i.e., the basic laws and principles of oceanic physics, biology, and chemistry, that evolve the state variables forward in time. Thus, predicting the present and future state of oceanic variables for practical applications is intimately linked to fundamental ocean science. A dynamical model to approximate nature consists of a set of coupled nonlinear prognostic field equations for each state variable of interest. The fundamental properties of the system appear in the field equations as parameters (e.g., viscosities, diffusivities, representations of body forces, rates of earth rotation, grazing, mortality, etc.). The initial and boundary

conditions necessary for integration of the equations may also be regarded as parameters by data assimilation methods. In principle the parameters of the system can be estimated directly from measurements. In practice, directly measuring the parameters of the interdisciplinary (physical-acoustical-optical-biological-chemical-sedimentological) ocean system is difficult because of sampling, technical, and resource requirements. However, data assimilation provides a powerful methodology for parameter estimation via the melding of data and dynamics. The physical state variables are usually the velocity components, pressure, density, temperature, and salinity. Examples of biological and chemical state variables are concentration fields of nutrients, plankton, dissolved and particulate matter, etc. Important complexities are associated with the vast range of phenomena, the multitude of concurrent and interactive scales in space and time, and the very large number of possible biological state variables. This complexity has two essential consequences. First, state variable definitions relevant to phenomena and scales of interest need to be developed from the basic definitions. Second, approximate dynamics which govern the evolution of the scale-restricted state variables, and their interaction with other scales, must be developed from the basic dynamical model equations. A familiar example consists of decomposing the basic ocean fields into slower and faster time scales, and shorter and longer space scales, and averaging over the faster and shorter scales. The resulting equations can be adapted to govern synoptic/mesoscale resolution state variables over a large-scale oceanic domain, with faster and smaller scale phenomena represented as parameterized fluctuation correlations (Reynolds stresses). There is, of course, a great variety of other scalerestricted state variables and approximate dynamics of vital interest in ocean science. We refer to scalerestricted state variables and approximate dynamics simply as ‘state variables’ and ‘dynamics’. The use of dynamics is of fundamental importance for efficient and accurate field and parameter estimation. Today and in the foreseeable future, data acquisition in the ocean is sufficiently difficult and costly so as to make field and parameter estimates by direct measurements, on a substantial and sustained basis, essentially prohibitive. However, data acquisition for field and parameter estimates via data assimilation is feasible, but substantial resources must be applied to obtain adequate observations.


1

2

DATA ASSIMILATION IN MODELS

The general process of state and parameter estimation is schematized in Figure 1. Measurement models link the state variables of the dynamical model to the sensor data. Dynamics interpolates and extrapolates the data. Dynamical linkages among state variables and parameters allow all of them to be estimated from measurements of some of them, i.e., those more accessible to existing techniques and prevailing conditions. Error estimation and error models play a crucial role. The data and dynamics are melded with weights inversely related to their relative errors. The final estimates should agree with the observations and measurements within data error bounds and should satisfy the dynamical model

Adaptive sampling

Observational network Sampling scheme Sensor data and data errors

UPl

Model improvements

LS

USVj

DE

MPk

M

OSVi

ION SCH E

_ DYNA M

DELS _ A MO SS L A

T ILA IM

IC

Measurement models and error models

_ ERROR M O ES

State and parameter estimates and errors SV SV: P: O: M: U: E: :

SVE

P

PE

State Variable Parameter Observed Measured Unobserved or Unmeasured Error Dynamical linkages

Figure 1 Data assimilation system schematic. Arrows represent the most common direction for the flows of information. The arrows between the measurement models and dynamical models double because measurement models can include operators that map state variables and parameters to the sensor data (e.g. interpolations, derivatives or integrals of state variables/parameters) and operators that transform sensor data into data appropriate for the model scales and processes (e.g. filtering, extrapolations or integrals of sensor data). The legend at the bottom explains abbreviations.

within model error bounds. Thus the melded estimate does not degrade the reliable information of the observational data, but rather enhances that information content. There are many important feedbacks in the generally nonlinear data assimilation system or ocean observing and prediction system (OOPS) schematized in Figure 1, which illustrates the system concept and two feedbacks. Prediction provides the opportunity for efficient sampling adapted to real time structures, events, and errors. Data collected for assimilation also used for ongoing verification can identify model deficiencies and lead to model improvements. A data assimilation system consists of three components: a set of observations, a dynamical model, and a data assimilation scheme or melding scheme. Modern interdisciplinary OOPS generally have compatible nested grids for both models and sampling. An efficient mix of platforms and sensors is selected for specific purposes. Central to the concept of data assimilation is the concept of errors, error estimation, and error modeling. The observations have errors arising from various sources: e.g., instrumental noise, environmental noise, sampling, and the interpretation of sensor measurements. All oceanic dynamical models are imperfect, with errors arising from the approximate explicit and parameterized dynamics and the discretization of continuum dynamics into a computational model. A rigorous quantitative establishment of the accuracy of the melded field and parameter estimates, or verification, is highly desirable but may be difficult to achieve because of the quantity and quality of the data required. Such verification involves all subcomponents: the dynamical model, the observational network, the associated error models, and the melding scheme. The concept of validation is the establishment of the general adequacy of the system and its components to deal with the phenomena of interest. As simple examples, a barotropic model should not be used to describe baroclinic phenomena, and data from an instrument whose threshold is higher than the accuracy of the required measurement are not suitable. In reality, validation issues can be much more subtle. Calibration involves the tuning of system parameters to the phenomena and regional characteristics of interest. Final verification requires dedicated experiments with oversampling. At this point it is useful to classify types of estimates with respect to the time interval of the data input to the estimate for time t. If only past and present data are utilized, the estimation is a filtering process. After the entire time series of data is available for (0,T), the estimate for any time 0 r t r T is



best based on the whole data set and the estimation is a smoothing process.

Goals and Purposes The specific uses of data assimilation depend upon the relative quality of data sets and models, and the desired purposes of the field and parameter estimates. These uses include the control of errors for state estimates, the estimation of parameters, the elucidation of real ocean dynamical processes, the design of experimental networks, and ocean monitoring and prediction. First consider ocean prediction for scientific and practical purposes, which is the analog of numerical weather prediction. In the best case scenario, the dynamical model correctly represents both the internal dynamical processes and the responses to external forcings. Also, the observational network provides initialization data of desired accuracy. The phenomenon of loss of predictability nonetheless inhibits accurate forecasts beyond the predictability limit for the region and system. This limit for the global atmosphere is 1–2 weeks and for the mid-ocean eddy field of the north-west Atlantic on the order of weeks to months. The phenomenon is associated with the nonlinear scale transfer and growth of initial errors. The early forecasts will accurately track the state of the real ocean, but longer forecasts, although representing plausible and realistic synoptical dynamical events, will not agree with contemporary nature. However, this predictability error can be controlled by the continual assimilation of data, and this is a major use of data assimilation today. Next, consider the case of a field estimate with adequate data but a somewhat deficient dynamical model. Assimilated data can compensate for the imperfect physics so as to provide estimates in agreement with nature. This is possible if dynamical model errors are treated adequately. For instance, if a barotropic model is considered perfect, and baroclinic real ocean data are assimilated, the field estimate will remain barotropic. Even though melded estimates with deficient models can be useful, it is of course important to attempt to correct the model dynamics. Parameter estimation via data assimilation is making an increasingly significant impact on ocean science via the determination of both internal and external parameter values. Regional field estimates can be substantially improved by boundary condition estimation. Biological modelers have been hampered by the inability to directly measure in situ rates, e.g., grazing and mortality. Thus, for interdisciplinary studies, internal parameter estimation is particularly promising. For example, measurements

3

of concentration fields of plankton together with a realistic interdisciplinary model can be used for in situ rate estimation. Data-driven simulations can provide four-dimensional time series of dynamically adjusted fields which are realistic. These fields, regarded as (numerical) experimental data, can thus serve as high resolution and complete data sets for dynamical studies. Balance of terms in dynamical equations and overall budgets can be carried out to determine fluxes and rates for energy, vorticity, productivity, grazing, carbon flux, etc. Case studies can be carried out, and statistics and general processes can be inferred for simulations of sufficient duration. Of particular importance are observation system simulation experiments (OSSEs), which first entered meteorology almost 30 years ago. By subsampling the simulated ‘true’ ocean, future experimental networks and monitoring arrays can be designed to provide efficient field estimates of requisite accuracies. Data assimilation and OSSEs develop the concepts of data, theory, and their relationship beyond those of the classical scientific methodology. For a period of almost 300 years, scientific methodology was powerfully established on the basis of two essential elements: experiments/observations and theory/ models. Today, due to powerful computers, science is based on three fundamental concepts: experiment, theory, and simulation. Since our best field and parameter estimates today are based on data assimilation, our very perception and conceptions of nature and reality require philosophical development. It is apparent from the above discussion that marine operations and ocean management must depend on data assimilation methods. Data-driven simulations should be coupled to multipurpose management models for risk assessments and for the design of operational procedures. Regional multiscale ocean prediction and monitoring systems, designed by OSSEs, are being established to provide ongoing nowcasts and forecasts with predictability error controlled by updating. Both simple and sophisticated versions of such systems are possible and relevant.

Regional Forecasting and Dynamics In this section, the issues and concepts introduced in the preceding sections are illustrated in the context of real-time predictions carried out in 1996 for NATO naval operations in the Strait of Sicily and for interdisciplinary multiscale research in 1998 in Massachusetts Bay. The Harvard Ocean Prediction System (HOPS) with its primitive equation dynamical model was utilized in both cases. In the Strait of Sicily (Figure 2), the observational network with platforms


4


Figure 2 Strait of Sicily. (A) Schematic of circulation features and dominant variabilities. (B) Forecast of the surface temperature for 25 August 1996, overlaid with surface velocity vectors (scale arrow is 0.25 m s1). (C) Objectively analyzed surface standard error deviation associated with the aircraft sampling on 18 September 1996 (normalized from 0 to 1). (D) Surface values of the first nondimensional temperature variability mode. (E) Satellite SST distributions for 25 August 1996. (F) Main LIW pathways, features, and mixing on deep potential density anomaly iso-surface (sY ¼ 29.05), over bottom topography (for more details, see Lermusiaux, 1999, and Lermusiaux and Robinson, 2001).



consisting of satellites, ships, aircraft, and Lagrangian drifters, was managed by the NATO SACLANT Undersea Research Centre. In Massachusetts Bay (Figure 3), the observational network with platforms consisting of ships, satellites, and autonomous underwater vehicles, was provided by the Littoral Ocean Observing and Prediction System (LOOPS) project within the US National Ocean Partnership Program. The data assimilation methods used in both cases were the HOPS OI and ESSE schemes (see Estimation Theory below). In both cases the purposes of data assimilation were to provide a predictive capability, to control loss of predictability, and to infer basic underlying dynamical processes. The dominant regional variabilities determined from these exercises and studies are schematized in Figures 2A and 3A. The dominant near surface flow in the strait is the Atlantic Ionian Stream, AIS (black lines for the stream; smooth and meandering dashed lines for the common locations of fronts and wave patterns, respectively) and dominant variabilities include the location and shapes of the Adventure Bank Vortex (ABV), Maltese Channel Crest (MCC), Ionian Shelfbreak Vortex (ISV), and Messian Rise Vortex (MRV) with shifts and deformations 0(10–100 km) occurring in 0(3–5 days). The variability of the Massachusetts Bay circulation is more dramatic. The buoyancy flow-through current which enters the Bay in the north from the Gulf of Maine may have one, two or three branches, and together with associated vortices (which may or may not be present), can reverse directions within the bay. Storm events shift the pattern of the features which persist inertially between storms. Actual real-time forecast fields are depicted in Figures 2B and 3B. The existence of forecasts allows adaptive sampling, i.e., sampling efficiently related to existing structures and events. Adaptive sampling can be determined subjectively by experience or objectively by a quantitative metric. The sampling pattern associated with the temperature objective analysis error map (Figure 2C) reflects the flight pattern of an aircraft dropping subsurface temperature probes (AXBTs). The data were assimilated into a forecast in support of naval operations centered near the ISV (Figure 2A). The sampling extends to the surrounding meanders of the AIS, which will affect the current’s thermal front in the operational region. The multiscale sampling of the Massachusetts Bay experiment is exemplified in Figure 3C by ship tracks adapted to the interactive submesoscales, mesoscales, bay-scales, and large-scales. Note that the tracks of Figure 3D are superimposed on a forecast of the total temperature forecast error standard deviation. The shorter track is objectively located around an error

5

maximum. The longer track is for reduction of velocity error (not shown). Eigendecomposition of variability fields helps dynamical interpretations. This eigen-decomposition estimates and orders the directions of largest variability variance (eigenmodes) and the corresponding amplitudes (eigenvalues). The first temperature variability eigenmodes for the strait and the bay are depicted in Figures 2D and 3E respectively. The former is associated with the dominant ABV variability and the latter with the location, direction, and strength of the inflow to the bay of the buoyancy current from the Gulf of Maine. A qualitative skill score for the prediction of dominant variations of the ABV, MCC, and ISV indicated correct 2- to 3-day predictions of surface temperature 75% of the time. The scores were obtained by validation against new data for assimilation and independent satellite sea surface temperature data as shown in Figure 2E for the forecast of Figure 2B. An important kinematical and dynamical interconnection between the eastern and western Mediterranean is the deep flow of salty Levantine Intermediate Water (LIW), which was not directly measured but was inferred from data assimilative simulations (Figure 2F). The scientific focus of the Massachusetts Bay experiment was plankton patchiness, in particular the spatial variability of zooplankton and its relationship to physical and phytoplankton variabilities (Figure 3B, G). The smallest scale measurements in the bay were turbulence measurements from an AUV (Figure 3F), which were also used to research the assimilation in real time of subgridscale data in the primitive equation model.

Concepts and Methods By definition (see Introduction), data assimilation in ocean sciences is an estimation problem for the ocean state, model parameters, or both. The schemes for solving this problem often relate to estimation or control theories (see below), but some approaches like direct minimization, stochastic, and hybrid methods (see below) can be used in both frameworks. Several schemes are theoretically optimal, while others are approximate or suboptimal. Although optimal schemes are preferred, suboptimal methods are generally the ones in operational use today. Most schemes are related in some fashion to least-squares criteria which have had great success. Other criteria, such as the maximum likelihood, minimax criterion or associated variations might be more appropriate when data are very noisy and sparse, and when probability density functions are multimodal (see Stochastic and hybrid models below). Parameters are assumed from here on to be


6


Figure 3 Massachusetts Bay. (A) Schematic of circulation features and dominant variabilities. (B) Chlorophyll-a (mg m3) at 10 m, with overlying velocity vectors. (C) Sampling pattern for the bay scales and external large-scales in the Gulf of Maine. (D) Forecast of the standard error deviation for the surface temperature (from 01C in dark blue to a maximum of 0.71C in red), with tracks for adaptive sampling. (E) 20 m values of the temperature component of the first nondimensional physical variability mode. (F) AUV turbulence data (Naval Underwater Warfare Center (NUWC)). (G) Vertical section of zooplankton (mM m3) along the entrance of Massachusetts Bay (for more details, see Robinson and the LOOPS group, 1999, and Lermusiaux, 2001).



included in the vector of state variables. For more detailed discussions, the reader is referred to the article published by Robinson et al. in 1998 (see Further Reading section).

7

data and forecast. Conditions for convergence to the Kalman filter have been derived, but in practice only a few iterations are usually performed. Frequently, the scales or processes of interest are corrected one after the other, e.g., large-scale first, then mesoscale.

Estimation Theory

Estimation theory computes the state of a system by combining all available reliable knowledge of the system including measurements and theoretical models. The a priori hypotheses and melding or estimation criterion are crucial since they determine the influence of dynamics and data onto the state estimate. At the heart of estimation theory is the Kalman filter, derived in 1960. It is the sequential, unbiased, minimum error variance estimate based upon a linear combination of all past measurements and dynamics. Its two steps are: (1) the forecast of the state vector and of its error covariance, and (2) the dataforecast melding and error update, which include the linear combination of the dynamical forecast with the difference between the data and model predicted values for those data (i.e., data residuals). The Kalman smoother uses the data available before and after the time of interest. The smoothing is often carried out by propagating the future data information backward in time, correcting an initial Kalman filter estimate using the error covariances and adjoint dynamical transition matrices, which is usually demanding on computational resources. In a large part because of the linear hypothesis and costs of these two optimal approaches, a series of approximate or suboptimal schemes have been employed for ocean applications. They are now described, from simple to complex. Direct insertion consists of replacing forecast values at all data points by the observed data which are assumed to be exact. The blending estimate is a scalar linear combination, with user-assigned weights, of the forecast and data values at all data points. The nudging or Newtonian relaxation scheme ‘relaxes’ the dynamical model towards the observations. The coefficients in the relaxation can vary in time but, to avoid disruptions, cannot be too large. They should be related to dynamical scales and a priori estimates of model and data errors. In optimal interpolation (OI), the matrix weighting the data residuals, or gain matrix, is empirically assigned. If the assigned OI gain is diagonal, OI and nudging schemes can be equivalent. However, the OI gain is usually not diagonal, but a function of empirical correlation and error matrices. The method of successive corrections performs multiple but simplified linear combination of the

Control Theory

All control theory or variational approaches perform a global time-space adjustment of the model solution to all observations and thus solve a smoothing problem. The goal is to minimize a cost function penalizing misfits between the data and ocean fields, with the constraints of the model equations and their parameters. The misfits are interpreted as part of the unknown controls of the ocean system. Similar to estimation theory, control theory results depend on a priori assumptions for the control weights. The dynamical model can be either considered as a strong or weak constraint. Strong constraints correspond to the choice of infinite weights for the model equations; the only free variables are the initial conditions, boundary conditions and/or model parameters. A rational choice for the cost function is important. A logical selection corresponds to dynamical model (data) weights inversely proportional to a priori specified model (data) errors. In an ‘adjoint method’, the dynamical model is a strong constraint. One penalty in the cost function weights the uncertainties in the initial conditions, boundary conditions, and parameters with their respective a priori error covariances. The other is the sum over time of data-model misfits, weighted by measurement error covariances. A classical approach to solve this constrained optimization is to use Lagrange multipliers. This yields Euler-Lagrange equations, one of which is the so-called adjoint equation. An iterative algorithm for solving these equations has often been termed the adjoint method. It consists of integrating the forward and adjoint equations successively. Minimization of the gradient of the cost function at the end of each iteration leads to new initial, boundary, and parameter values. Another iteration can then be started, and so on, until the gradient is small enough. Expanding classic inverse problems to the weak constraint fit of both data and dynamics leads to generalized inverse problems. The cost function is usually as in adjoint methods, except that a third term now consists of dynamical model uncertainties weighted by a priori model error covariances. In the Euler-Lagrange equations, the dynamical model uncertainties thus couple the state evolution with the adjoint evolution. The representer method is an algorithm for solving such problems.


8


Direct Minimization Methods

Examples

Such methods directly minimize cost functions similar to those of generalized inverse problems, but often without using the Euler-Lagrange equations. Descent methods iteratively determine directions locally ‘descending’ along the cost function surface. At each iteration, a minimization is performed along the current direction and a new direction is found. Classic methods to do so are the steepest descent, conjugate-gradient, Newton, and quasi-Newton methods. A drawback for descent methods is that they are initialization sensitive. For sufficiently nonlinear cost functions, they are restarted to avoid local minima. Simulated annealing schemes are based on an analogy to the way slowly cooling solids arrange themselves into a perfect crystal, with a minimum global energy. To simulate this relatively random process, a sequence of states is generated such that new states with lower energy (lower cost) are always accepted, while new states with higher energy (higher cost) are accepted with a certain probability. Genetic algorithms are based upon searches generated in analogy to the genetic evolution of natural organisms. They evolve a population of solutions mimicking genetic transformations such that the likelihood of producing better data-fitted generations increases. Genetic algorithms allow nonlocal searches, but convergence to the global minimum is not assured due to the lack of theoretical base.

This section presents a series of recent results that serve as a small but representative sample of the wide range of research carried out as data assimilation was established in physical oceanography.

Stochastic and Hybrid Methods

Stochastic methods are based on nonlinear stochastic dynamical models and stochastic optimal control. Instead of using brute force like descent algorithms, they try to solve the conditional probability density equation associated with ocean models. Minimum error variance, maximum likelihood or minimax estimates can then be determined from this probability density. No assumptions are required, but for large systems, parallel machines are usually employed to carry out Monte Carlo ensemble calculations. Hybrid methods are combinations of previously discussed schemes, for both state and parameter estimation; for example, error subspace statistical estimation (ESSE) schemes. The main assumption of such schemes is that the error space in most ocean applications can be efficiently reduced to its essential components. Smoothing problems based on Kalman ideas, but with nonlinear stochastic models and using Monte Carlo calculations, can then be solved. Combinations of variational and direct minimization methods are other examples of hybrid schemes.

General Circulation from Inverse Methods

The central idea is to combine the equations governing the oceanic motion and relevant oceanic tracers with all available noisy observations, so as to estimate the large-scale steady-state total velocities and related internal properties and their respective errors. The work of Martel and Wunsch in 1993 exemplifies the problem. The three-dimensional circulation of the North Atlantic (Figure 4A) was studied for the period 1980–85. The observations available consisted of objective analyses of temperature, salinity, oxygen, and nutrients data; climatological ocean–atmosphere fluxes of heat, water vapor, and momentum; climatological river runoffs; and current meter and float records. These data were obtained with various sensors and platforms, on various resolutions, as illustrated by Figure 4B. A set of steady-state equations were assumed to hold a priori, up to small unknown noise terms. The tracers were advected and diffused. The advection velocities were assumed in geostrophic thermal–wind balance, except in the top layer where Ekman transport was added. Hydrostatic balance and mass continuity were assumed. The problem is inverse because the tracers and thermal wind velocities are known; the unknowns are the fields of reference level velocities, vertical velocity, and tracer mixing coefficients. Discrete finite-difference equations were integrated over a set of nested grids of increasing resolutions (Figure 4A). The flows and fluxes at the boundaries of these ocean subdivisions were computed from the data (at 11 resolution). The resulting discrete system contained c. 29 000 unknowns and 9000 equations. It was solved using a tapered (normalized) least-squares method with a sparse conjugate-gradient algorithm. The estimates of the total flow field and of its standard error are plotted on Figure 1C and D. The Gulf Stream, several recirculation cells and the Labrador current are present. In 1993, such a rigorous large-scale, dense, and eclectic inversion was an important achievement. Global versus Local Data Assimilation via Nudging

Malanotte-Rizzoli and Young in 1994 investigated the effectiveness of various data sets to correct and control errors. They used two data sets of different types and resolutions in time and space in the Gulf



9

60˚N

84 85 54˚N 48˚N 42˚N 36˚N 30˚N

32 27

18˚N

28 10˚W

18˚W

12˚N 27˚W

26 36˚W

25 53˚W

62˚W

(B)

GM

(A)

GM

71˚W

24

24˚N

45˚W

31

2.

5

_ 60

2.5

2.0

2.0

0

2.5

3.0

30

30 50

equator

equator

(C)

(D)

Figure 4 (A) Domain of the model used in the inverse computations. Weak dynamical constraints were imposed on the flow and tracers, and integrated over a set of nested grids, from the full domain (heavy solid lines) to an ensemble of successive divisions (e.g., dashed lines) reaching at the smallest scales the size of the boxes labeled by numbers. (B) Locations of the stations where the hydrographic and chemical component of the data set were collected (model grid superposed). (C) Inverse estimate of the absolute sea surface topography in centimeters (contour interval is 10 cm). (D) Inverse estimate of the standard error deviation (in cm) of the sea surface topography shown in (C). (Reproduced with permission from Martel and Wunsch, 1993.).

Stream region, at mesoscale resolution and for periods of the order of 3 months, over a large-scale domain referred to as global scale. One objective was to assimilate data of high quality, but collected at localized mooring arrays, and to investigate the effectiveness of such data in improving the realistic attributes of the simulated ocean fields. If successful, such estimates allow for dynamical and process studies. The global data consisted of biweekly fields of sea surface dynamic

height, and of temperature and salinity in three dimensions, over the entire region, as provided by the Optimal Thermal Interpolation Scheme (OTIS) of the US Navy Fleet Numerical Oceanography Center. The local data were daily current velocities from two mooring arrays. The dynamical model consisted of primitive equations (Rutgers), with a suboptimal nudging scheme for the assimilation. The global and local data were first assimilated alone, and then together. The ‘gentle’ assimilation of


10


Cross-over points

. = Representers calculated

(A)

60

TPX 0.1: M 2 30

0

_ 30

_ 60

0

(B)

60

0

120

30

180

60

300

240

90

120

360

150

cm

Figure 5 (A) TOPEX/POSEIDON crossover points and subsampling. Representers were calculated only for the windowed subset of 986 satellite crossover points (large filled dots), but differences from all 6355 crossover points (small dots and large filled dots) were included in the data-misfit penalty. (B) Generalized inverse estimate of the amplitude and phase of the M2 tidal constituent. The phase isolines are plotted in white over color-filled contours of the amplitude. Contour interval is 10 cm for amplitude and 301 for phase. (Reproduced with permission from Egbert et al. 1994.).

the spatially dense global OTIS data was necessary for the model to remain on track during the 3-month period. The ‘strong’ assimilation of the daily but local data from the Synoptic Ocean Prediction SYNOP was required to achieve local dynamical accuracies, especially for the velocities. Small-scale Convection and Data Assimilation

In 1994, Miller et al. addressed the use of variational or control theory approaches (see earlier) to assimilate data into dynamical models with highly

nonlinear convection. Because of limited data and computer requirements, most practical ocean models cannot resolve motions that result from static instabilities of the water column; these motions and effects are therefore parameterized. A common parameterization is the so-called convective adjustment. This consists of assigning infinite mixing coefficients (e.g., heat and salt conductivities) to the water at a given level that has higher density than the water just below. This is carried out by setting the densities of the two parcels to a unique value in such a way that heat and mass are conserved. In a



numerical model, at every time step, water points are checked and all statically unstable profiles replaced by stable ones. The main issues of using such convective schemes with variational data assimilation are that: (1) the dynamics is no longer governed by smooth equations, which often prevents the simple definition of adjoint equations; (2) the optimal ocean fields may evolve through ‘nonphysical’ states of static instability; and (3), the optimization is nonlinear, even if the dynamics are linear. Ideally, the optimal fields should be statically stable. This introduces a set of inequality constraints to satisfy. An idealized problem was studied so as to provide guidance for realistic situations. A simple variational formulation had several minima and at times produced evolutions with unphysical behavior. Modifications that led to more meaningful solutions and suggestions for algorithms for realistic models were discussed. One option is the ‘weak’ static stability constraint: a penalty that ensures approximate static stability is added to the cost function with a very small error or large weight. In that case, static stability can be violated, but in a limited fashion. Another option is the ‘strong constraint’ form of static stability which can be enforced via Lagrange multipliers. Convex programming methods which explicitly account for inequality constraints could also be utilized.

11

size of the problem was reduced by winnowing out the full set of 6350 crossovers to an evenly spaced subset of 986 points (see Figure 5A). The resulting representer matrix was then reduced by singular value decomposition. The amplitude and phase estimates for the M2 constituent are shown in Figure 5B. The M2 fields are qualitatively similar to previous results and amphidromes are consistent. However, when compared with previous tidal model estimates, the inversion result is noticeably smoother and in better agreement with altimetric and ground truth data.

Conclusions The melding of data and dynamics is a powerful, novel, and versatile methodology for parameter and field estimation. Data assimilation must be anticipated both to accelerate research progress in complex, modern multiscale interdisciplinary ocean science, and to enable marine technology and maritime operations that would otherwise not be possible.

Acknowledgments We thank Ms. G. Sweetland and Ms. M. Armstrong for help in preparing the manuscript, and the ONR for partial support.

Global Ocean Tides Estimated from Generalized Inverse Methods

In 1994, Egbert et al. estimated global ocean tides using a generalized inverse scheme with the intent of removing these tides from the data collected by the TOPEX/POSEIDON satellite and thus allowing the study of subtidal ocean dynamics. A scheme for the inversion of the satellite crossover data for multiple tidal constituents was applied to 38 cycles of the data, leading to global estimates of the four principal tidal constituents (M2, S2, K1 and O1) at about 11 resolution. The dynamical model was the linearized, barotropic shallow water equations, corrected for the effects of ocean self-attraction and tidal loading, the state variables being the horizontal velocity and sea surface height fields. The data sets were linked to measurement models and comprehensive error models were derived. The generalized inverse tidal problem was solved by the representer method. Representer functions are related to Green’s functions: they link a given datum to all values of the state variables over the period considered. These representers were computed by solving the Euler-Lagrange equations in parallel. The

See also Open Ocean Convection. Biogeochemical Data Assimilation. Coastal Circulation Models. Current Systems in the Mediterranean Sea. CTD (Conductivity, Temperature, Depth) Profiler. Elemental Distribution: Overview. Expendable Sensors. Florida Current, Gulf Stream and Labrador Current. Forward Problem in Numerical Models. Inverse Models. Mesoscale Eddies. Ocean Circulation. Patch Dynamics. Regional and Shelf Sea Models. Tides. Upper Ocean Time and Space Variability.

Further Reading Anderson D and Willebrand J (eds.) (1989) Oceanic Circulation ModelsC: ombining Data and Dynamics. Dordrecht: Kluwer Academic. Bennett AF (1992) Inverse Methods in Physical Oceanography. Cambridge Monographs on Mechanics and Applied Mathematics. Cambridge: Cambridge University Press.


12


Brasseur P and Nihoul JCJ (eds.) (1994) Data assimilationt: ools for modelling the ocean in a global change perspective. Series IG: lobal Environmental Change, 19. Berlin: Springer-Verlag. NATO ASI Series. Egbert GD, Bennett AF, and Foreman MGG (1994) TOPEX/POSEIDON tides estimated using a global inverse model. Journal of Geophysical Research 24: 821--824 852. Haidvogel DB and Robinson AR (eds) (1989) Special issue on data assimilation. Dynamics of Atmospheres and Oceans 13: 171–517. Lermusiaux PFJ (1999) Estimation and study of mesoscale variability in the Strait of Sicily. Dynamics of Atmospheres and Oceans 29: 255--303. Lermusiaux PFJ (2001) Evolving the subspace of the threedimensional multiscale ocean variability: Massachusetts Bay. Journal of Marine Systems. In press. Lermusiaux PFJ and Robinson AR (1999) Data assimilation via error subspace statistical estimation, Part IT: heory and schemes. Monthly Weather Review 7: 1385--1407. Lermusiaux PFJ and Robinson AR (2001) Features of dominant mesoscale variability, circulation patterns and dynamics in the Strait of Sicily. Deep Sea Research, Part I. In press. Malanotte-Rizzoli P and Young RE (1995) Assimilation of global versus local data sets into a regional model of the Gulf Stream system: I. Data effectiveness. Journal of Geophysical Research 24: 773--796. Malanotte-Rizzoli P (ed.) (1996) Modern Approaches to Data Assimilation in Ocean Modeling, Elsevier

Oceanography Series. The Netherlands: Elsevier Science. Martel F and Wunsch C (1993) The North Atlantic circulation in the early 1980s – an estimate from inversion of a finite-difference model. Journal of Physical Oceanography 23: 898--924. Miller RN, Zaron EO, and Bennett AF (1994) Data assimilation in models with convective adjustment. Monthly Weather Review 122: 2607--2613. Robinson AR, Lermusiaux PFJ, and Sloan NQ III (1998) Data assimilation. In: Brink KH and Robinson AR (eds.) The SeaT: he Global Coastal Ocean I, Processes and Methods, 10. New York: John Wiley and Sons. Robinson AR (1999) Forecasting and simulating coastal ocean processes and variabilities with the Harvard Ocean Prediction System. In: Mooers CNK (ed.) Coastal Ocean Prediction, AGU Coastal and Estuarine Study Series, pp. 77--100. Washington: AGU Press. Robinson AR and the LOOPS group (1999) Real-time forecasting of the multidisciplinary coastal ocean with the Littoral Ocean Observing and Predicting System (LOOPS). Third Conference on Coastal Atomspheric and Oceanic Prediction and Processes. New Orleans, LA: American Meterological Society, 30*35. (3-5 Nov 1999). Robinson AR and Sellschopp J (2000) Rapid assessment of the coastal ocean environment. In: Pinardi N and Woods JD (eds.) Ocean ForecastingC: onceptual Basis and Applications. London: Springer-Verlag. Wunsch C (1996) The Ocean Circulation Inverse Problem. Cambridge: Cambridge University Press.


DEEP CONVECTION J. R. N. Lazier, Bedford Institute of Oceanography, NS, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 634–643, & 2001, Elsevier Ltd.

Introduction Density of ocean water generally increases with depth except at the surface where stirring by waves and convection creates a well-mixed homogeneous layer. Breaking waves alone can mix the upper 5–10 m, but convection, forced by an increase in density at the surface via heat loss or evaporation, can greatly increase the mixed layer depth. During winter, heat loss from the surface of the ocean is high and convectively mixed surface layers are the norm in the extratropical oceans. The deepest (4 1500 m) are found in the Labrador Sea, the Greenland Sea, and the Golfe du Lion in the Mediterranean Sea, because of two special features. First, they are near land where cold continental air flows over the water to create the necessary high heat loss. Second, the circulation in each is weakly cyclonic which helps to maintain the convecting water where the high heat loss occurs. The combination of these features provides the persistent heat loss from the same body of water that is needed to force convection to reach great depths. The example of a deepening convection layer in Figure 1 shows profiles of s1.5 (s1.5 þ 1000 ¼ potential density in kg m 3 referenced to 1500 decibars; 1 decibar corresponds to about 1 m) versus depth, on February 25 and March 8, 1997, in the Labrador Sea. The convecting layer is the approximately homogenous layer next to the surface about 750 m deep with a s1.5 of 34.65 kg m 3 on February 25 (Station 1) and 1150 m deep and 34.67 kg m 3 11 days later (Station 2). The buoyancy that was removed between the two profiles is proportional to the area between them, i.e. buoyancy ¼ ðg=r0 Þ

Z Drdz

where g (10 m s 2) is the acceleration due to gravity, r0(E1034 kg m 3) is the reference density, z is the depth and Dr is the difference in density between Stations 1 and 2. For the two stations illustrated this calculation yields a buoyancy loss of 0.17 m2 s 2. By

ignoring the small effect of evaporation, precipitation, and any advection, this buoyancy loss can be assumed to be due solely to heat loss from the surface. The loss is converted to joules by dividing by ga/r0c where g and r0 are as before and a (about 10 4 1C 1) is the thermal expansion of water and c (4.2 kJ kg 1 1C 1) is the specific heat of sea water at constant pressure. The conversion suggests that a heat loss of about 0.68 109 J m 2 was required to remove the buoyancy between the two dates. Over the 11 days between the observations this heat loss is equivalent to an average rate of heat loss of 715 W m 2. By a similar calculation, the heat loss required to increase the depth of convection to 2000 m would have been about 1.2 109 J m 2 or 460 W m 2 over a month. If the profiles in Figure 1 had been obtained during an era of mild winters rather than during one of abnormally severe winters they probably would have exhibited a markedly lower density in the upper layers. This might occur because of abnormally large freshwater flows into the surface layers due to increased outflows from the Arctic, warmer summers or a multiyear period of restratification following a vigorous period of convection. As the lower density represents ‘extra’ buoyancy to be removed before convection can proceed to greater depths, the ultimate depth of the convecting layer will be less in this situation, for a given heat loss, than in the illustrated one. Thus the ultimate depth of the convecting layer during a winter depends on the total amount of buoyancy lost from the sea surface and the distribution of that buoyancy with depth. The distribution of the newly convected water in two dimensions is illustrated in the contour plot of salinity across the Labrador Sea in Figure 2. These data were obtained in July following the exceptionally cold winter of 1992–93. The water mass resulting from convection is the large volume of nearly homogeneous water lying between 360 and 800 km on the horizontal scale and between 500 and 2300 m in the vertical. Because of its large volume, unique properties, and the fact that it spreads beyond its region of formation, this water is known as Labrador Sea Water. The upper layer (0–500 m) in the central part of the section is clearly not as well mixed as the layer between 500 and 2300 m. This is because the observations were obtained in July about 3 months after deep convection ceased at the end of the cooling season, about April 1. Since that time the surface layer has been flooded with fresh water derived from melting ice and river runoff. Also, the layer below this


13

14

DEEP CONVECTION

1.5 (kg m–3) 34.64 0

34.66

34.68

Sta 1

34.70

34.72

Sta 2

400

Depth (m)

800

1200

1600

2000

Figure 1 Vertical distribution of s1.5 obtained from R/V Knorr at 56.81N, 54.21W in the Labrador Sea on February 25 (Station 1) and March 8, 1997 (Station 2).

1

5

10

15

20

25

0 34.93 34.92 34.91 34.90 34.89 34.88 34.87 34.86 34.85 34.84 34.83 34.82 34.81 34.80 34.75 34.70 34.60 34.40 34.20 34.00 33.50 33.00 32.00

Depth (m)

1000

2000

3000

100

300

500

700

900

Distance (km) Figure 2 Salinity distribution across the Labrador Sea between 53.01N, 55.51W and 60.61N, 49.31W obtained between June 19 and 23, 1993. The water between 500 m and 2200 m in the central part of the section is unusually homogeneous because of deep convective mixing during the severe winter of 1992–93. The CTD station positions are indicated by numbered triangles along the surface.


DEEP CONVECTION

low salinity surface layer, to about 500 m, has been invaded by higher salinity water from the right (northeast) and to a lesser extent the left (south-west). The newly formed Labrador Sea Water is a mixture of all the water down to 2300 m including the water in contact with the atmosphere at the surface. In these uppermost layers the concentration of gases such as oxygen, carbon dioxide, tritium, and chlorofluorocarbons (CFCs) are at or near equilibrium with the atmosphere. By transporting these gases down from the upper layers of the ocean to intermediate depths, convection provides a mechanism to ventilate the deeper layers, which is one of the most important consequences of deep convection. Dissolved oxygen, for example, is slowly used up in the deep ocean by biological processes and would eventually vanish without the renewal via convection. Also most of the carbon dioxide ever put in the atmosphere by volcanoes since the formation of the Earth became dissolved in the ocean and is now contained in sediments in the bottom of the ocean. As combustion of fossil fuels over the earth raises the carbon dioxide content of the atmosphere it is important to understand the rate at which this gas is entering the deeper layers of the ocean through processes such as deep convection. Subsequent to formation, Labrador Sea Water spreads to other regions of the ocean at intermediate depths. Knowledge of the speed of this flow and its influence increased during the 1990s due to the widespread high quality observations of temperature, salinity, and CFCs across the North Atlantic obtained under the international World Ocean Circulation Experiment. The newly ventilated water formed in the Labrador Sea during the severe winters of the early 1990s moved across the North Atlantic at about 2 cm s 1. This speed is about three to four times greater than the previous estimate, leading to the conclusion that the intermediate flows are much faster than previously thought. A comparison of six decades of data from the Labrador Sea and from the subtropical waters near Bermuda suggest that the products of deep convection in the Labrador Sea impact the waters off Bermuda after about 6 years.

Plumes – the Mixing Agent Convection begins to increase the depth of the mixed layer in the Labrador Sea near the end of September when the surface net buoyancy flux from the surface turns from positive to negative. Deepening continues until about the end of March when the buoyancy flux again becomes positive. When convection is active, water at the surface becomes denser than the underlying water and descends in plumes. This water

15

is replaced by slightly lighter water rising toward the surface. The physical features of the convecting water including the plumes and the water between have been the subject of a number of investigations; most notably by the group of scientists at Kiel working in the Golfe du Lion in the Mediterranean Sea with moored acoustic Doppler current profilers (ADCP) and current meters. The cartoon in Figure 3 summarizes some of the main features of plumes and the mixing layer. At the surface is the thermal boundary layer where the water is losing heat/buoyancy to the atmosphere. Water in this layer is, on average, slightly denser than in the mixed layer beneath and descends into the mixing layer within plumes which have a horizontal dimension of about 1 km, approximately equal to that vertical extent, i.e. an aspect ratio of E 1. The average rate of descent within the plumes is about 0.02 m s 1 while the maximum is E0.13 m s 1. Rotation of the plumes, due to the horizontal component of the Coriolis force, is expected because water must converge into the plume at its top and presumably diverge out of it near the bottom. However, this effect has not yet been conclusively observed in the field although it has been observed in laboratory experiments and in numerical simulations. Another effect that has not been observed is an increasing horizontal dimension with depth which is expected if water is entrained into the plumes as they descend, or if they decelerate as they go deeper. Observations also indicate that there is no net vertical mass flux within a convecting region or patch. This appears to have solved the long-standing puzzle of whether the descending water was replaced by rising water between the plumes or by converging flow in the upper layer and diverging flow in the deep layer. Finally, on average, the plumes are not penetrative, i.e. the plumes do not have enough energy to descend into water that is denser than the water within the plume. One consequence of this is illustrated in Figure 1 by the fact that the bottom of the mixing layer at Station 2 lies on the s1.5 versus depth curve observed earlier on February 25. If convection was penetrative the bottom of the mixed layer on March 8 would lie below this curve and the s1.5 versus depth gradient at the bottom of the mixing layer would be greater than when it was observed earlier. The plumes in the cartoon suggest that there should be a high correlation between fluctuations in temperature and vertical velocity. However, recent measurements from drifting floats indicate this correlation to be weak, thus making the cartoon a rather simplified view of the true situation. In reality the plumes are probably not vertical over the full depth


16

DEEP CONVECTION

Figure 3 A schematic diagram of a E 1000 m convecting layer indicating approximate values for features of individual plumes including the horizontal scale, vertical downward velocity, rotation, and entrainment. Across the patch of convecting water there is no net vertical mass flux and no significant penetration into the layers of denser water beneath the convecting layer. Buoyancy (B0) lost from the surface creates the thermal boundary layer in the upper E 100 m where the denser water is formed which sinks within the plumes. The wiggly up arrow indicates the (slow) upward flow that replaces the (relatively fast) downward flow within the plumes. Note that the horizontal scale in the figure is E 50 times the vertical scale.

of the convecting layer but contorted by the largerscale flows. A direct view of the motion within convecting plumes has recently been obtained from freely drifting floats. When one of these is launched it immediately sinks to a predetermined depth below the convecting layer where it remains for typically 7 days while its buoyancy adjusts. At the end of this period its buoyancy is decreased slightly and the float rises into the convecting layer. A large attached drogue then causes the float to moved up and down with the vertical motions of convection. In the Labrador Sea over 25 days in February and March 1997 the maximum vertical velocity observed by a set of these floats was downward at 0.2 m s 1 with a rms value for all the observations of 0.02 m s 1. This is equivalent to a round trip of the convecting layer of 1 day for the average water parcel. On a number of

occasions the floats were seen to penetrate below the average bottom of the mixing layer. Contrary to the conclusions mentioned above that the convection is not penetrative, this suggests that a certain amount of plume penetration into denser layers does occur.

Temperature and Salinity Variability Year-long time-series records of temperature and salinity obtained in the middle of the Labrador Sea indicate that there is a marked increase in temperature and salinity variability during and following convection. This is evident in the temperature record in Figure 4 obtained at 510 m in 1994–95. Between June and the middle of February the temperature sensor is below the mixed layer and the temperature slowly increases by about 0.21C. In mid-February the temperature drops by about 0.41C as the deepening


DEEP CONVECTION

17

3.4

Potential temperature (°C)

3.2

3.0

2.8

2.6

2.4 June

Aug

Oct

Dec

Feb

Apr

June

Months, 1994 – 95 Figure 4 A 1 year record of temperature at 510 m in the central Labrador Sea illustrating the increase in variance during and after the mixed layer reaches the depth of the instrument in early February 1995 plus the sudden upward shift in temperature near April 1 associated with the end of convection.

convecting layer reaches the depth of the sensor. At this time the magnitude of the variations in temperature suddenly increase and continue at a high level for a number of weeks. Spectra calculated from 85-day pieces of this record before and after the arrival of the convection layer show a broadband fourfold increase in the spectral energy of the variability after the sensor is immersed in the mixed layer. Time-series of the energy show a peak in February shortly after the mixed layer arrives followed by a decline to 1/30th of the peak value by the end of the record in June. Similar fluctuations occur in salinity and are largely in phase with those in temperature. Figure 5 shows temperature versus salinity plots at four depths during 21 days in March 1995 when convection was proceeding to about E 2000 m. During this time period density at each of these levels was relatively constant in time. In the figure the hourly observations at each depth show the extent of the fluctuations in temperature and salinity and the fact that they are largely parallel to the constant density surfaces; the T–S fluctuations tend to be parallel to the isopycnals. The most probable explanation for the fluctuations is that they are horizontal variations in temperature and salinity being swept past the mooring by the current. One suggestion is that these variations reflect horizontal variability in the depth of convection. However, recent work in the mixed

layer of the tropical Pacific demonstrates that compensating horizontal variations in temperature and salinity may be ubiquitous features of the mixed layer whether it is convecting or not. These results lead to the alternative suggestion that the compensating temperature and salinity variations exist in the windmixed surface layer before deep convection begins and are propagated downward by the convection. When convection stops, the vertical density stratification is reestablished in the surface layers and the horizontal variations that appeared during convection are no longer renewed. Those variations existing when convection ends are then slowly mixed away by turbulent eddies leading to the decay in the amplitude of the fluctuations as observed in the timeseries. Assuming a horizontal scale L of 100 km for the region of convection with its small-scale horizontal variations, the timescale of eddy mixing will be about L2/KH where KH, the horizontal eddy diffusivity, is about 103 m2 s 1. This gives a timescale for the horizontal mixing of about 4 months, which is about the decay time observed in the records.

Restratification At the end of the cooling season, vertical mixing due to convection ceases and its dominant influence on mid-depth water properties ends. This also marks the beginning of the restratification process during which


18

DEEP CONVECTION

260 m 3.0

510 m

3.0

2

4.6


2.9

1.5

=3

2.9

2.8

2.8

2.7

2.7 2.6

2.6 .72

2.5 34.78

34

34.80

34.82

34.84

2.5 34.86

34.78

34.80


3.0

2.9

2.9

2.8

2.8

2.7

2.7

2.6

2.6

34.80

34.82

34.84

34.86

34.84

34.86

1510 m

1010 m 3.0

2.5 34.78

34.82

34.84

34.86

2.5 34.78

Salinity

34.80

34.82 Salinity

Figure 5 Temperature versus salinity diagrams based on data obtained at 260, 510, 1010, and 1510 m in the central Labrador Sea, between March 12 and April 2, 1995, during the final stage of convection when the density at these depths remained relatively constant.

the vertical stratification existing prior to the homogenization begins to be reestablished. Two timescales seem to be involved. At the end of convection a rapid restratification occurs which is indicated by sudden shifts of variables such as temperature. One example is indicated in Figure 4 by the increase of 0.21C near April 1 following roughly 6 weeks of convective activity at this depth. It is not clear if this increase indicates an end to convection over a large area, or the advection of a stratified nonconvecting water column to the observation site. The first option, however, seems more likely as the end of convection appears in other records as a rapid increase in stratification especially in the upper layers. For example, a tomographic array in the Golfe du Lion observed a roughly 40 day restratification period following convection. This seems to be the only observation of restratification over a large area. Another example is the record from a PALACE

(profiling autonomous Lagrangian circulation explorer) in the Labrador Sea which shows, during 2 consecutive years, a sudden transition between the low stratification associated with convection and a stratified water column. This record is admittedly like a mooring, from a single point, but it does give a consistent picture in the 2 years. A recent numerical model of the restratification process may describe this rapid phase. It has a homogeneous cylinder of water floating in an ocean of constant stratification. The density gradient between the cylinder and the surrounding ocean gives rise to a narrow cyclonic current which breaks up via baroclinic instability into baroclinic eddies. These mix the homogeneous water horizontally with the stratified waters and so dissipate the homogeneous cylinder in timescale t. For the special case where the stratification in the water surrounding the homogeneous cylinder is concentrated in the upper


DEEP CONVECTION

layer h:

19

Restratification over a number of years is illustrated in the time series of s1.5 through the 1990s in Figure 7. In the early years of the decade the water produced by the convection, indicated by the layer of low vertical gradient, lies within the 34.64 and 34.70 kg m 3 surfaces. Between 1990 and 1994 the volume of this waterremains roughly constant but the value of s1.5 at the core increases from 34.67 to 34.69 kg m 3 as the winters became more severe and the convecting layer continued to deepen into the stratified layer below. In the years following 1995, convection was limited to 1000–1500 m. The deep reservoir of ‘homogeneous’ water was thereby isolated below the convecting layer and slowly decreased in volume with each passing year as it drained away. This decrease in volume was balanced by an increase in the volume of lighter more stratified water in the upper layers from the boundaries. The large interannual variation in the volume of the Labrador Sea Water illustrated in this figure is a well known property of the water mass; however, its effects on the large-scale ocean currents and processes are not yet well understood.

tE56 r=ðhDbÞ1=2 where r is the radius of the homogeneous cylinder, h is the depth of the upper stratified water bounding the homogeneous cylinder and Db is the difference in density between the homogeneous water and the surrounding water in buoyancy units (i.e. Db ¼ gDr/r0). For the Labrador Sea where h E 500 m, Db E 2 10 3 m s 2 and r E 100 km; t E 65 days. This result, being of the same order as the observed rapid changes in stratification, suggests that the model may be appropriate to explain the observations. The long timescale of restratification has now been observed in the Labrador Sea. Observations have been made continuously over one summer and intermittently over a few successive years. Changes over the summer of 1996 are illustrated in Figure 6 by the depths of four of the isopycnals within the upper 1000 m. At stations between 400 and 600 km the isopycnal depths increase significantly between May and October while the 27.72 and 27.74 kg m 3, surfaces between 650 and 790 km, show a decrease in depth. These changes suggest that lighter water, from beyond the region of deepest convection (400– 600 km), moves into the upper water column in the region of deepest convection while denser water in the region of deepest convection moves outward toward the boundaries at mid-depth.

Discussion While the focus of this article has been on the Labrador Sea, numerous aspects of the convective processes discussed here also apply to the other two locations of open-ocean convection in the North

Distance (km) 200

300

400

500

600

700

800

900

0 27.64 200

Depth (m)

27.68

400 27.72 May 600 October 27.74 800

1000 Figure 6 Four surfaces of constant s0 (potential density anomaly relative to 0 m) across the Labrador Sea based on CTD data collected in May 1996 (solid lines) and October 1996 (dotted lines).


20

DEEP CONVECTION

0

1.5 = 34.56

500 34.60

Depth (m)

1000 34.64 1500

34.68

2000

2500 1990

1992

1994

1996

1998

2000

Years Figure 7 Depths of s1.5 surfaces in the middle of the Labrador Sea through the 1990s based on CTD data collected in May or June of each year.

Atlantic: the Greenland Sea and the Mediterranean Sea. But as there are common threads to the overturning in these seas there are also significant regional differences. The Greenland Sea is unique in that ice plays a significant role in the preconditioning phase. The deepest convective overturning occurs in late winter just after the ice-free ‘Nord Bukta’ region opens up. In the Mediterranean the winds are more localized than in the other two regions. This clearly influences the convection, as the region of deepest mixed layers generally lies in the path of the Mistral winds. Furthermore, these winds are rarely coldenough to cause convection during daylight hours, so the Mediterranean has a daily cycle of overturning that is not present in the other two seas. Also, the impact of the basin-scale NorthAtlantic Oscillation wind pattern influences the Labrador and Greenland Seas to a much greater extent than the Mediterranean. Finally, the dimensions of the convection zones and convected water masses differ greatly. In the Mediterranean, the convecting patch is of order 50 km wide; in the Greenland Sea it is of order 100 km wide, and in the Labrador Sea the zone of convection approaches 500 km in width, and includes the boundary currents. Convection at each of these three Atlantic sites contributes to the global meridional overturn circulation, although the quantitative measures are not yet known. In the Greenland Sea particularly, the deep convection into the cyclonic gyre seems rather isolated from the processes that produce the dense

overflows. Mediterranean Water and Labrador Sea Water both make an obvious contribution to the Upper North Atlantic Deep Water, respectively, as high and low salinity endpoints. Distant identification of Labrador Sea Water is through its low salinity; potential vorticity; low nutrient concentration; high dissolved oxygen, tritium, and CFCs. Along the western boundary velocity and CFC maxima associated with Labrador Sea Water have been observed near Abaco, nearly 5000 km south. For the era ending in 1977 a dilution of the tritium maxima of the deep western boundary currents by factors of order 10 was observed, from the subpolar gyre to the Blake-Bahama Outer Ridge. This suggested dilution and delay (recirculation) mechanisms en route. Model studies suggest that when convection is initiated or increased at the high-latitude source, a pressure wave propagates south along the western boundary, as a topographic Rossby wave, well before the arrival of tracer-tainted, identifiable water mass. Such model studies point out that sloping topography acts as a wave guide, and rather gently leads dense water masses equatorward from high latitude, as they slowly sink. Thus, ‘sinking’ is minimal in the near-field of the convection, but occurs downstream. Production of kinetic energy of theoverturning circulation by potential energy created by buoyancy forcing requires that dense water sink and less dense water rise, but the sites of sinking and rising are, at least in modestudies, often distant from the


DEEP CONVECTION

convection. In a diapycnal/epipycnal coordinate system, however, convection is more locally associated with time-averaged diapycnal transport (‘watermass conversion’). Such analyses are beginning to be carried out with models and are an insightful way to approach the link between convection, sinking, and global meridional overturning. Thus we are still seeking to quantify production rates of the constituent water masses of the global meridional overturning. Outward transport of Labrador Sea Water is even difficult todefine, because of extensive recirculation within the subpolar gyre, and entrainment once the water mass has left the subpolar gyre. Estimates have ranged from o1 Sv to 410 Sv. Much alsremains unknown about the detailed geography of deep convection and circulation. In the Labrador Sea, both interior and boundary currents are known to participate in the deep convection (estimates of 1–2 Sv of boundary current production). However, the boundary current, is shielded from deep convection by low salinity shelf waters at some sites. Where the circulation crosses from Greenland to Labrador, the boundary currents broaden and slow down, andare generally exposed to some of the most intense air–sea heat flux in the sea; there and over the wide continental slope near Labrador, convection may be particularly deep. Direct velocity and transport measurements are needed to augment water mass observations. Unfortunately the Lagrangian movement of water masses is difficult to observe even with modern ‘quasi-Lagrangian’ floats and drifters. A recent description of the Labrador/Irminger Sea circulation from PALACE floats notes that ‘no floats travelled southward to the subtropical gyre in the deep western boundary current, the putative main pathway of dense

21

water in the meridional overturning circulation’. If the boundary current is concentrated to a narrow width, for example at the Flemish Cap, then these profiling floats may have difficulty staying within it; tracer observations assure us that the transport does in fact take place.

See also Carbon Cycle. Current Systems in the Mediterranean Sea. Mediterranean Sea Circulation. Rossby Waves. Sub-sea Permafrost.

Further Reading Lazier JR, Pickart RS, and Rhines PB (2001) Deep convection. In: Ocean Circulation and Climate – Observing and Modelling the Global Ocean. London: Academic Press. Lilly J, Rhines P, Visbeck M, et al. (1999) Observing deep convection in the Labrador Sea during winter 1994– 1995. Journal of Physical Oceanography 29: 2065--2098. Marshall J and Schott F (1999) Open-ocean convection observations, theory and models. Reviews of Geophysics 37: 1--64. Schott R, Visbeck M, and Send U (1994) Open ocean deep convection, Mediterranean and Greenland Seas. In: Malanotte-Rizzoli P and Robinson AR (eds.) Ocean Processes on Climate Dynamics: Global and Mediterranean Examples, pp. 203--225. Dordrecht: Kluwer Academic Publishers. Lab Sea Group (1998) The Labrador Sea Deep Convection Experiment. Bulletin of the American Meteorological Society 79: 2033--2058.


DEEP SUBMERGENCE, SCIENCE OF D. J. Fornari, Woods Hole Oceanographic Institution, Woods Hole, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 643–658, & 2001, Elsevier Ltd.

Introduction The past half-century of oceanographic research has demonstrated that the oceans and seafloor hold the keys to understanding many of the processes responsible for shaping our planet. The Earth’s ocean floor contains the most accurate (and complete) record of geologic and tectonic history for the past 200 million years that is available for a planet in our solar system. For the past 30 years, the exploration and study of seafloor terrain throughout the world’s oceans using ship-based survey systems and deep submergence platforms has resulted in unraveling plate boundary processes within the paradigm of sea floor spreading; this research has revolutionized the Earth and Oceanographic sciences. This new view of how the Earth works has provided a quantitative context for mineral exploration, land utilization, and earthquake hazard assessment, and provided conceptual models which planetary scientists have used to understand the structure and morphology of other planets in our solar system. Much of this new knowledge stems from studying the seafloor – its morphology, geophysical structure, and characteristics, and the chemical composition of rocks collected from the ocean floor. Similarly, the discoveries in the late 1970s of deep sea ‘black smoker’ hydrothermal vents at the midocean ridge (MOR) crest (Figure 1) and the chemosyntheticbased animal communities that inhabit the vents have changed the biological sciences, provided a quantitative context for understanding global ocean chemical balances, and suggest modern analogs for the origin of life on Earth and extraterrestrial life processes. Intimately tied to these research themes is the study of the physical oceanography of the global ocean water masses and their chemistry and dynamics, which has resulted in unprecedented perspectives on the processes which drive climate and climate change on our planet. These are but a few of the many examples of how deep submergence research has revolutionized our understanding of our Earth and ocean history, and provide a glimpse at the diversity of scientific frontiers that await exploration in the years to come.

22

Enabling Deep Submergence Technologies The events that enabled these breakthroughs was the intensive exploration that typified oceanographic expeditions in the 1950s to 1970s, and focused development of oceanographic technology and instrumentation that facilitated discoveries on many disciplinary levels. Significant among the enabling technologies were satellite communication and global positioning, microchip technology and the widespread development of computers that could be taken into the field, and increasingly sophisticated geophysical and acoustic modeling and imaging techniques. The other key enabling technologies which supplanted traditional mid-twentieth century methods for imaging and sampling the seafloor from the beach to the abyss were submersible vehicles of various types, remote-sensing instruments, and sophisticated acoustic systems designed to resolve a wide spatial and temporal range of ocean floor and oceanographic processes. Oceanographic science is by nature multidisciplinary. Science carried out using deep submergence vehicles of all types has traditionally involved a wide range of research components because of the time and expense involved with conducting field research on the seafloor using human occupied submersibles or remotely operated vehicles (ROVs), and most recently autonomous underwater vehicles (AUVs) (Figures 2–6). Through the use of all these vehicle systems, and most recently with the advent of seafloor observatories, deep submergence science is poised to enter a new millennium where scientists will gain a more detailed understanding of the complex linkages between physical, chemical, biological, and geological processes occurring at and beneath the seafloor in various tectonic settings. Understanding the temporal dimension of seafloor and sub-seafloor processes will require continued use of deep ocean submersibles and utilization of newly developed ROVs and AUVs for conducting timeseries and observatory-based research in the deep ocean and at the seafloor. These approaches will provide new insights into intriguing problems concerning the interrelated processes of crustal generation, evolution, and transport of geochemical fluids in the crust and into the oceans, and origins and proliferation of life both on Earth and beyond. Since the early twentieth century, people have been venturing into the ocean in a wide range of diving vehicles from bathyscapes to deep diving submersibles.


DEEP SUBMERGENCE, SCIENCE OF

(A)

23

(B)

(D)

(C)

Figure 1 Hydrothermal vents on the southern East Pacific Rise axis at depths of 2500–2800 m. (A) Titanium fluid sampling bottles aligned along the front rail of Alvin’s basket in preparation for fluid sampling. (B) Alvin’s manipulator claw preparing to sample the hydrothermal sulfide chimney. (C) View from inside Alvin’s forward-looking view port of the temperature probe being inserted into a vent orifice to measure the fluid temperature. (D) Hydrothermal vent after a small chimney was sampled which opened up the orifice through which the fluids are exiting the seafloor. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group, D. J. Fornari, K. Von Damm, and M. Lilley.

Even in ancient times, there was written and graphic evidence of the human spirit seeking the mysteries of the ocean and seafloor. There is unquestionably the continuing need to take the unique human visual and cognitive abilities into the ocean and to the seafloor to make observations and facilitate measurements. For about the past 40 years submersible vehicles of various types have been developed largely to support strategic naval operations of various countries. As a result of that effort, the US deep-diving submersible Alvin was constructed. Alvin is part of the National Deep Submergence Facility (NDSF) of the University National Oceanographic Laboratories System (UNOLS) operated by the Woods Hole Oceanographic Institution (Figure 7). Alvin provides routine scientific and engineering access to depths as great as 4500 m. The

US academic research community also has routine, observational access to the deep ocean and seafloor down to 6000 m depth using the ROV and tethered vehicles of the NDSF (Figure 7). These vehicle systems of the NDSF include the ROV Jason, and the tethered optical/acoustic mapping systems Argo II and DSL120 sonar (a 120 kHz split-beam sonar system capable of providing 1–2 m pixel resolution back-scatter imagery of the seafloor and phase-bathymetric maps with B4 m pixel resolution (Figure 7). These fiberopticbased ROV and mapping systems can work at depths as great as 6000 m. Alvin has completed over 3600 dives (more than any other submersible of its type), and has participated in making key discoveries such as: imaging, mapping, and sampling the volcanic seafloor on


24


(A)

(A)

(B)

Figure 3 (A) The ROV Jason being lifted off the stern of a research vessel at the start of a dive. (B) ROV Jason recovering amphora on the floor of the Mediterranean Sea. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group and R. Ballard.

(B)

Figure 2 (A) The submersible Alvin being lifted onto the stern of R/V Atlantis, its support ship. (B) Alvin descending to the seafloor. Photos courtesy of Woods Hole Oceanographic Institution – Alvin Group and R. Catanach.

the MOR crest; structural, petrological, and geochemical studies of transforms faults; structural studies of portions of deep-sea trenches off Central America; petrological and geochemical studies of volcanoes in back-arc basins in the western Pacific Ocean; sedimentary and structural studies of submarine canyons, discovering MOR hydrothermal vents; and collecting samples and making time series measurements of biological communities at hydrothermal vents in many MOR settings in the Atlantic

and Pacific Oceans. In 1991, scientists in Alvin were also the first to witness the vast biological repercussions of submarine eruptions at the MOR axis, which provided the first hint that a vast subsurface biosphere exists in the crust of the Earth on the ocean floor (Figure 7).

Deep Submergence Science Topics Some of the recent achievements in various fields of deep submergence science include the following. 1. Discoveries of deep ocean hydrothermal communities and hot (43501C) metal-rich vents on



(A)

(D)

(B)

(E)

(C)

(F)

25

Figure 4 Photographs taken from the ROV Jason at hydrothermal vents on the Mid-Atlantic Ridge near 371N on the summit of Lucky Strike Seamount at a depth of 1700 m. All photographs are of a vent named ‘Marker d4.’ (A) Overall view of vent looking North. White areas are anhydrite and barite deposits, yellowish areas are covered with clumps of vent mussels. (B) Close-up of the side of the vent; Jason’s sampling basket is in the foreground. (C) Close-up view of mussels on the side of the hydrothermal chimney; the hydrothermal vent where hot fluids are exiting the mound is at the upper right, where the image is blurry because of the shimmering effect of the hot water. Nozzles of a titanium sampling bottle are at the middle-right edge. (D) Inserting a self-recording temperature probe into a beehive chimney. (E) Titanium fluid sampling bottles being held by Jason’s manipulator during sampling. (F) Close-up of nozzles of sampling bottle inserted into vent orifice during sampling. Photos (D)–(F) are frame grabs of Jason video data. Photos courtesy of Woods Hole Oceanographic Institution – ROV Group, D. J. Fornari, S. Humphris, and T. Shank.


26


(A)

(B)

Figure 5 (A) ROV Tiburon of the Monterey Bay Aquarium Research Institute (MBARI) in the hanger of its support ship R/V Western Flyer. This electric-powered ROV can dive to 4000 m. A steel-armored, electro-optical cable connects Tiburon to the R/V Western Flyer and delivers power to the vehicle. Electric thrusters allow fine maneuvering while minimizing underwater noise and vehicle disturbance. A variable buoyancy control system, together with the syntactic foam pack, enables Tiburon to hover inches above the seafloor without creating turbulence, to pick up a rock sample, or maneuver quickly to follow an animal. (B) ROV Ventana of the MBARI being launched from its support ship R/V Pt. Lobos. This ROV gives researchers the opportunity to make remote observations of the seafloor to depths of 1850 m. The vehicle has two manipulator arms – a seven-function arm with five spatially correspondent joints and another sevenfunction robot arm with six spatially correspondent joints. Both arms can use a variety of end effectors to suit the type of work being done. Ventana is also equipped with a conductivity, temperature, and density (CTD) package including a dissolvedoxygen sensor and a transmissometer. Photos courtesy of MBARI.

many segments of the global mid-ocean ridge (MOR); 2. Documentation of the immediate after-effects of submarine eruptions on the northern East Pacific Rise, Axial Seamount, Gorda Ridge and CoAxial Segment of the Juan de Fuca Ridge;

Figure 6 The autonomous underwater vehicle (AUV) ABE (Autonomous Benthic Explorer) of the Woods Hole Oceanographic Institution’s Deep Submergence Laboratory, which can survey the seafloor completely autonomously to depths up to 4000 m and is especially well suited to working in rugged terrain such as is found on the Mid Ocean Ridge. Photo courtesy of Woods Hole Oceanographic Institution.

3. Utilization of Ocean Drilling Program bore holes and specialized vehicle systems (e.g. Scripps Institution’s Re-Entry Vehicle) (Figure 8) and instrument suites (e.g. CORKs) (Figure 9) for a wide range of physical properties, fluid flow and seismological experiments; 4. Discoveries of extensive fluid flow and vent-based biological communities along continental margins and subduction zones; 5. Initial deployment of ocean floor observatories of various types which enable the monitoring and sampling of geological, physical, biological and chemical processes at and beneath the seafloor (Figures 10 and 11). These studies have revolutionized our concepts of deep ocean processes and highlighted the need



27

Figure 7 (A) Summary of the US National Deep Submergence Facility (NDSF) vehicles operated for the University National Oceanographic Laboratories System (UNOLS) by the Woods Hole Oceanographic Institution (WHOI). (B) Montage of the NDSF vehicles showing examples of the various types of data they collect. The figure also shows the nested quality of the surveys conducted by the various vehicles which allows scientists to explore and map features with dimensions of tens of kilometers (top right multibeam sonar map), to detailed sonar back scatter and bathymetry swaths which have a pixel resolution of 1–2 m (DSL-120 sonar), which are then further explored with the Argo II imaging system, or sampled using Alvin or ROV Jason. Graphics by P. Oberlander, WHOI; photos courtesy of WHOI – Alvin and ROV Group.


28


_ 1500 _ 2000 _ 2500

37˚ 05′ N

_ 3000 _ 3500

32˚ 10′ W 32˚ 20′ W 10 m

Depth (m)

_ 1000 37˚ 25′ N

DSL 120 Sonar

N

Multibeam Sonar Argo II Imaging System

1 km DSV Alvin

3m

ROV Jason

0.75 m

5m (B)

Figure 7 (Continued)

for more detailed, time-series, multidisciplinary research. Within the field of biological oceanography, major recent advances have come from the study of the new life forms and chemoautotrophic

processes discovered at hydrothermal vents (Figure 10). These advances have fundamentally altered biological classification schemes, extended the known thermal and chemical limits of life, and have pushed the search for origins of life on Earth



(A)

Figure 8 (A) The Scripps Institution of Oceanography’s Control Vehicle is a specialized ROV that can place instrument strings inside deep-sea boreholes, using a conventional oceanographic vessel capable of dynamic positioning and equipped with a winch carrying 17.3 mm (0.68 in) electromechanical cable with a single coax (RG8-type). The control vehicle is 3.5 m tall and weighs about 500 kg in water (1000 kg in air). It consists of a stainlesssteel frame that contains two orthogonal horizontal hydraulic thrusters, a compass, a Paroscientific pressure gauge, four 250 W lights, a video camera, sonar systems, electronic interfaces to electrical releases and to a logging probe, and electronics to control all these sensors and handle data telemetry to and from the ship. Telemetry on the tow cable’s single coax is achieved by analog frequency division multiplexing over a frequency band extending from 20 kHz to about 800 kHz. The sonars include a 325 kHz sector-scanning sonar, a 23.5 kHz narrow beam acoustic altimeter, and a 12 kHz sonar for long baseline acoustic navigation. This system was used successfully in the 1998 Ocean Seismic Network Pilot Experiment at ODP Hole 843B, 225 km south-west of the island of Oahu, Hawaii. In 1999, the control vehicle’s analog telemetry module was converted into a digital system using fiberoptic technology thus providing bandwidth capabilities in excess of 100 Mbaud. (B) Cartoon showing the configuration of the control vehicle deployed from a research ship as it enters an Ocean Drilling Program bore hole to insert an instrument string. Photo and drawing courtesy of Scripps Institution of Oceanography – Marine Physical Laboratory, F. Spiess, and C. de Moustier.

29

as well as for new life forms on other planetary bodies. Recent marine biological studies show that: (1) the biodiversity of every marine community is vastly greater than previously recognized; (2) both sampling statistics and molecular tools indicate that the large majority of marine species have not been described; (3) the complexity of biological communities is far greater than previously realized; and (4) the response of various communities to both natural and anthropogenic forcing is far more complex than had been understood even a decade ago. Focused studies over long time periods will be required to characterize these communities fully. The field of marine biology is heading toward a more global time-series approach as a function of recent discoveries largely in the photic zone of the oceans and in the deep ocean at MOR hydrothermal vents. The marine biological, chemical, and physical oceanographic research which will be carried out in the next decade and beyond will certainly have a profound impact on our understanding of the complex food webs in the ocean which control productivity at every level and have direct implications for commercial harvesting of a wide range of resources from the ocean. Meeting the challenge of deciphering the various chemical, biological, and physical influences on these phenomena will require a better understanding and resolution of the causes and consequences of change on scales from hours to millennia. Understanding ocean ecosystems and their constituents will improve dramatically in response to emerging molecular, chemical, optical, and acoustical technologies. Given the relative paucity of information on deep-sea fauna in general, and especially the relatively recent discovery of chemosynthetic ecosystems at MOR hydrothermal vents, this will continue to be a focus for deep ocean biological research in the coming decade and beyond. Time-series and observatory-based research and sampling techniques will be required to answer the myriad of questions regarding the evolution and physiology of these unique biological systems (Figures 10 and 11). Present and future foci for deep submergence science is MOR crests, hydrothermal systems, and the volcano–tectonic processes that create the architecture of the Earth’s crust. Geochemists from the marine geology and geophysics community have emphasized the need for studies of: (1) the flux-frequency distribution for ridge-crest hydrothermal activity (heat, fluid, chemistry); (2) the role played by fluid flow in gas hydrate accumulation and determination of how important hydrates are to climate change; (3) slope


30


Soft tether and Multiconductor cable 0.68″ Control vehicle

Tow cable

300 m

Probe

Acoustic transponder

Cone

(B) Figure 8b (Continued)

stability; and (4) determination of how much of a role the microbial community plays in subsurface chemical and physical transformations. Another focus involves subduction zone processes, including: an assessment of the fluxes of fluids and solids through the seismogenic zone; long-term monitoring of changes in seismicity, strain, and fluid flux in the seismogenic zone; and determining the nature of materials in the seismogenic zone. To answer these types of questions, the marine geology and geophysics community requires systematic studies of temporal evolution of diverse areas

on the ocean floor through research that includes mapping, dating, sampling, geophysical investigations, and drilling arrays of crustal holes (Figures 8, 9 and 11). The researchers in this field stress that the creation of true seafloor observatories at sites with different tectonic variables, with continuous monitoring of geological, hydrothermal, chemical, and biological activity will be necessary. Whereas traditional geological and geophysical tools will continue to provide some means to address aspects of these problems, it is clear that an array of deep submergence vehicles, in situ sensors, and ocean floor



31

Figure 9 (A) Diagram of the upper portion of an Ocean Drilling Program (ODP) borehole with a Circulation Obviation Retrofit Kit (CORK) assembly. These units serve the same purpose as a ‘cork’ which seals a bottle; in the deep-sea case, the bottle is the sea floor which contains fluids that are circulating in the ocean crust. The CORK allows scientists to access the circulating fluids and make controlled hydrologic measurements of the pressures and physical properties of the fluids. (B) A CORK observatory on the ocean floor in ODP hole 858G off the Pacific north-west coast. (C) A CORK with instruments installed to measure sub-seafloor fluid circulation processes. Diagram courtesy of Woods Hole Oceanographic Institution and J. Doucette; photo courtesy of K. Becker and E. Davis.

observatory systems will be required to address these topics and unravel the variations in the processes that occur over short (seconds/minutes) to decadal timescales. The infrastructural requirements, facility, and

development needs required to support the research questions to be asked include: a capability for longterm seafloor monitoring; effective detection and response capability for a variety of seafloor events


32


(A)

(B)

(F)

(C)

(D) (G)

(E) Figure 10 Time-series sequence of photographs taken of the same area of seafloor from the submersible Alvin of a hydrothermal vent site on the East Pacific Rise axis near 9149.80 N at a depth of 2500 m. (A) ‘Snow blower’ vent spewing white bacterial by-product during the 1991 eruption. (B) Same field of view as (A) about 9 months later. Diffuse venting is still occurring as is bacterial production. White areas in the crevices

of the lava flow are juvenile tube worms. (C) Patches of Riftia tube worms colonizing the vent area B18 months after the March 1991 eruption. (D) Tube worm community has continued to develop, and the venting continues over 5 years after the eruption. (E) Close-up photograph of zoarcid vent fish (middleleft), tube worms, mussels (yellowish oblong individuals) and bathyurid crab (center). (F) A time-series temperature probe (with black and yellow tape), deployed at a hydrothermal vent (Tube worm pillar) on the East Pacific Rise crest near 9149.60 N at a depth of 2495 m. The vent is surrounded by a large community of tube worms. (G) Seafloor markers along the Bio-Geo Transect, a series of 210 markers placed on the seafloor in 1992 to monitor the changes in hydrothermal vent biology and seafloor geology over a 1.4 km long section of the East Pacific Rise axis that have occurred after the 1991 volcanic eruption at this site. Photographs courtesy of T. Shank, Woods Hole Oceanographic Institution and R. Lutz, Rutgers University, D. J. Fornari and Woods Hole Oceanographic Institution – Alvin Group.



33

(A) Figure 11 (A) Diagram of the deployment of the Hawaii-2 Observatory (H2O) one of the first long-term, deep seafloor observatories deployed in the past few years. Scientists used the ROV Jason to splice an abandoned submarine telephone cable into a termination frame which acts as an undersea telephone jack. Attached by an umbilical is a junction box, which serves as an electrical outlet for up to six scientific instruments. An ocean bottom seismometer and hydrophone are now functioning at this observatory. (B) The H2O junction box as deployed on the seafloor and photographed by ROV Jason. Drawing courtesy of Jayne Doucette, Woods Hole Oceanographic Institution (WHOI), A. Chave, and WHOI–ROV group.

(volcanic, seismic, chemical); adequately supported, state-of-the-art seafloor sampling and observational facilities (e.g. submersible, ROVs, and AUVs), and accurate navigation systems, software, and support for shipboard integration of data from mutiscalar and nested surveys (Figures 7B, 11 and 12). As discussed above, the disciplines involved in deep-submergence science are varied and the scales of investigation range many orders of magnitude from molecules and micrometer-sized bacteria to segment-scales of the MOR system (10s to 100s of kilometers long) at depths that range from 1500 m to

6000 m and greater in the deepest trenches. Clearly, the spectrum of scientific problems and environments where they must be investigated require access to the deep ocean floor with a range of safe, reliable, multifaceted, high-resolution vehicles, sensors, and samplers, operated from support ships that have global reach and good station-keeping capabilities in rough weather. Providing the right complement of deep-submergence vehicles and versatile support ships from which they can operate, and the funding to operate those facilities cost-effectively, is both a requirement and a challenge for satisfying the


34


(B) Figure 11 (Continued).

objectives of deep-sea research in the coming decade and into the twenty-first century. To meet present and future research and engineering objectives, particularly with a multidisciplinary approach, deep submergence science will require a mix of vehicle systems and infrastructures. As deep submergence science investigations extend into previously unexplored portions of the global seafloor, it is critical that scientists have access to sufficient vehicles with the capability to sample, observe and make timeseries measurements in these environments. Submersibles, which provide the cognitive presence of humans and heavy payload capabilities will be critical to future observational, time-series and observatorybased research in the coming decades. Fiberopticbased ROVs and tethered systems, especially when used in closely timed, nested investigations offer unparalleled maneuverability, mapping and sampling capabilities with long bottom times and without the limitation of human/vehicle endurance. AUVs of various designs will provide unprecedented access to the global ocean, deep ocean and seafloor without dedicated support from a surface ship. AUVs represent vanguard technology that will revolutionize seafloor and oceanographic measurements and observations in the decades to come. Over approximately the past 5 years scientists at several universities and private laboratories have made enormous advances in the capabilities and fieldreadiness of AUV systems. One such system is the Autonomous Benthic Explorer (ABE) developed by engineers and scientists at the Woods Hole Oceanographic Institution (Figure 6). ABE can survey the seafloor completely autonomously and is especially well suited to working in rugged terrain such as is found on the MOR. ABE maps the seafloor and the water column near the bottom without any guidance

from human operators. It follows programmed track-lines precisely and follows the bottom at heights from 5 to 30 m, depending on the type of survey conducted. ABE’s unusual shape allows it to maintain control over a wide range of speeds. Although ABE spends most of its survey time driving forward at constant speed, ABE can slow down or even stop to avoid hitting the seafloor. In practice, ABE has surveyed areas in and around steep scarps and cliffs, and not only survived encounters with the extreme terrain, but also obtained good sensor data throughout the mission. Recently ABE has been used for several geological and geophysical research programs on the MOR in the north-east and south Pacific which have further proved its reliability as a seafloor survey vehicle, and pointed to its unique characteristics to collect detailed, near-bottom geological and geophysical data, and to ground-truth a wide range of seafloor terrains (Figure 12). These new perspectives on seafloor geology and insights into the geophysical properties of the ocean crust have greatly improved our ability to image the deep ocean and seafloor and have already fostered a paradigm shift in field techniques and measurements which will surely result in new perspectives for Earth and oceanographic processes in the coming decades.

Conclusions One of the most outstanding scientific revelations of the twentieth century is the realization that ocean processes and the creation of the Earth’s crust within the oceans may determine the livability of our planet in terms of climate, resources, and hazards. Our discoveries may even enable us to determine how life itself began on Earth and whether it exists on other worlds. The next step is toward discovering the linkages between various phenomena and processes in the oceans and in exploring the interdependencies of these through time. Marine scientists recognize that technological advances in oceanographic sensors and vehicle capabilities are escalating at a increasingly rapid pace, and have created enormous opportunities to achieve a scope of understanding unprecedented even a decade ago. This new knowledge will build on the discoveries in marine sciences over the last several decades, many of which have been made possible only through advances in vehicle and sensor technology. With the rapidly escalating advances in technology, marine scientists agree that the time is ripe to focus efforts on understanding the connections in terms of interdependency of phenomena at work in the world oceans and their variability through time.



_ _ 2 _ 2 46 4 24 5 0 40 0

46˚31 ′30″N

_ 24

30

46˚31 ′30″N

46˚31 ′00″N

_ 247

0

_ 24

90

46˚31 ′00″N

Magnetic Field (nT) 59500 59000 58500 58000 57500 57000 56500 56000 55500 55000 54500 54000 53500 53000 52500 52000 51500 51000 50500 50000 49500 49000 48500 48000

35

46˚30 ′30″N (A)

46˚30 ′30″N (B)

Thickness (meters) 45 24 22 20 18 16 14 12 10 8 6 4 2 1 0

46˚31 ′30″N

46˚31 ′00″N

400m

46˚30 ′30″N (C)

129˚35 ′30″N

46˚31 ′30″N

46˚31 ′00″N

400m

129˚35 ′00"N

129˚34 ′30″N

129˚35 ′30″N (D)

129˚35 ′00″N

129˚34 ′30″N

46˚30 ′30″N

Figure 12 (A) Bathymetry map (contour interval 10 m) showing location of the 1993 CoAxial lava eruption (gray) on the Juan de Fuca Ridge off the coast of Washington, and ABE tracklines (each color is a separate dive). (B) Magnetic field map based on ABE tracklines showing strong magnetic field over new lava flow. (C) Computed lava flow thickness assuming an average lava magnetization of 60 A/m compared with (D) lava flow thickness determined from differential swath bathymetry. Figure courtesy of M. Tivey, Woods Hole Oceanographic Institution.


36


See also Cephalopods. Deep-Sea Ridges, Microbiology. Hydrothermal Vent Fluids, Chemistry of. Manned Submersibles, Deep Water. Mid-Ocean Ridge Geochemistry and Petrology. Platforms: Autonomous Underwater Vehicles. Remotely Operated Vehicles (ROVs). Seamounts and Off-Ridge Volcanism.

Further Reading Becker K and Davis EE (2000) Plugging the seafloor with CORKs. Oceanus 42(1): 14--16. Chadwick WW, Embley RW, and Fox C (1995) Seabeam depth changes associated with recent lava flows, coaxial segment, Juan de Fuca Ridge: evidence for multiple eruptions between 1981–1993. Geophysical Research Letters 22: 167--170. Chave AD, Duennebier F, and Butler R (2000) Putting H2O in the ocean. Oceanus 42(1): 6--9. de Moustier C, Spiess FN, Jabson D, et al. (2000) Deep-sea borehole re-entry with fiber optic wire line technology. Proceedings 2000 of the International Symposium on Underwater Technology. pp. 23–26. Embley RW and Baker E (1999) Interdisciplinary group explores seafloor eruption with remotely operated vehicle. Eos Transactions of the American Geophysical Union 80(19): 213--219 222. Fornari DJ, Shank T, Von Damm KL, et al. (1998) Timeseries temperature measurements at high-temperature hydrothermal vents, East Pacific Rise 91490 –510 N: monitoring a crustal cracking event. Earth and Planetary Science Letters 160: 419--431.

Haymon RH, Fornari DJ, Von Damm KL, et al. (1993) Direct submersible observation of a volcanic eruption on the Mid-Ocean Ridge: 1991 eruption of the East Pacific Rise crest at 91450 –520 N. Earth and Planetary Science Letters 119: 85--101. Humphris SE, Zierenberg RA, Mullineaux L, and Thomson R (1995) Seafloor Hydrothermal systems: Physical, Chemical, Biological, and Geological Interactions. American Geophysical Union Monograph, vol. 91, 466 pp. Ryan WBF (chair) et al., (Committee on seafloor observatories: challenges and opportunities) (2000) Illuminating the Hidden Planet, the Future of Seafloor Observatory Science. Washington, DC: Ocean Studies Board, National Research Council, National Academy Press. Shank TM, Fornari DJ, Von Damm KL, et al. (1998) Temporal and spatial patterns of biological community development at nascent deep-sea hydrothermal vents along the East Pacific Rise, 9149.60 N–9150.40 N. Deep Sea Research, II 45: 465--515. Tivey MA, Johnson HP, Bradley A, and Yoerger D (1998) Thickness measurements of submarine lava flows determined from near-bottom magnetic field mapping by autonomous underwater vehicle. Geophysical Research Letters 25: 805--808. UNOLS (University National Laboratory System) (1994) The Global Abyss: An Assessment of Deep Submergence Science in the United States. Narragansett, RI: UNOLS Office, University of Rhode Island. Von Damm KL (2000) Chemistry of hydrothermal vent fluids from 9–101N, East Pacific Rise: ‘Time zero’ the immediate post-eruptive period. Journal of Geophysical Research 105: 11203--11222.


DEEP-SEA DRILLING METHODOLOGY K. Moran, University of Rhode Island, Narragansett, RI, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction The technology developed and used in past scientific drilling programs, the Deep Sea Drilling Project and the Ocean Drilling Program, has now been expanded in the current Integrated Ocean Drilling Program (IODP). These technologies include innovative drilling methods, sampling tools and procedures, in situ measurement tools, and seafloor observatories. This new IODP technology will be used to drill deeper into the seafloor than was possible in the previous scientific programs. The first drilling target of the IODP using the new technology is the seismogenic zone offshore Japan, a location deep in the Earth (7–14 km) where earthquakes are generated.

ship’s motion from the drill pipe; and a pump that flushes sea water through the drill pipe. Open hole methods are successfully used in all of the Earth’s oceans (Figure 2). Scientific ocean-drilling achievements include drilling in very deep water (6 km) and to 42 km below the seafloor (Table 1). Although there have been many achievements using these methods, there are also limitations. Although the exact depth limit of open-hole drilling method is not yet known, it is likely limited to 2–4 km below the seafloor. This limitation exists because the drill fluid must be modified to a lower density so that the deep cuttings can be lifted from the bit and flushed out of the hole when drilling deep into the seafloor. Another

Drilling Technology The Deep Sea Drilling Project and the Ocean Drilling Program used the same basic drilling technology, the open hole method. Today, the IODP has extended this capability to include closed hole methods, known as riser drilling. Open Hole or Nonriser Drilling

Drilling is the process of establishing a borehole. The open hole method uses a single drill pipe that hangs from the drill ship’s derrick, a tall framework positioned over the drill hole used to support the drill pipe. The drill pipe is rotated using drilling systems, specifically a hydraulically powered top drive located above the drill floor of the ship. Surface sea water is flushed through the center of the pipe to lubricate the rotating bit that cuts the rock and then flushes sediment and rock cuttings away to the seafloor (Figure 1). Open hole refers to the resulting borehole which remains open to the ocean during drilling. This method is also called a riserless drilling system. Important parts of the deep-water drilling system are a drilling derrick that is large and strong enough to hang a long length of drill pipe reaching deep ocean and subseafloor depths (up to 8 km); a system that rotates the drill pipe; a motion compensator that isolates the

Drill pipe

Drilling fluid is pumped down through drill pipe

Cuttings

Drilling fluid and cuttings flow into ocean

Seafloor Surface casing

Second casing

Uncased hole

Drilling fluid and cuttings flow up between the drill pipe and the borehole or casing

Drill bit

Figure 1 Diagram of a nonriser drilling system.


37


Figure 2 Map of all sites drilled by the Deep Sea Drilling Project (DSDP), the Ocean Drilling Program (ODP), and the IODP.

DEEP-SEA DRILLING METHODOLOGY

Table 1

39

Fact sheet about the JOIDES Resolution, the research vessel used by the Ocean Drilling Program

Total number of days in port Total number of days at sea Total distance traveled Total number of holes drilled Deepest water level drilled Deepest hole drilled

Leg 129 Leg 148, Hole 504B (South-eastern Pacific Ocean, off coast of Ecuador)

Total amount of core Most core recovered on single Northern-most site drilled

Leg 175 – Benguela (15 Aug.–10 Oct. 1997) Leg 113, site 911

Southern-most site drilled

Leg 151, site 693

Year and place of constitution

1978

Laboratories and other scientific equipment installed Gross tonnage Net tonnage Engines/generators

1984

Length Beam Derrick Speed Crusing range Scientific and technical party Ship’s Crew Laboratory space Drill string

limitation of the open hole method is that drilling must be restricted to locations where hydrocarbons are unlikely to be encountered. In an open hole, there is no way to control the drilling fluid pressure. In locations where oil and gas may exist, the formations are frequently overpressured (similar to a champagne bottle). If these formations would be punctured with an open hole system, the drill pipe would act like a straw that connects this overpressured zone in the rock to the ocean and the ship. This type of puncture is called a ‘blow-out’ and is a serious drilling hazard. The explosion as gases are vented through the straw to the ship’s drill floor could cause serious damage, or worse yet, the change in the density of the seawater as the gas bubbles are released into the overlying ocean could cause the ship to sink. Without a system to control the pressure in the borehole, there is no way to prevent a blow-out. Closed Hole or Riser Drilling

The new deep-sea drilling technology used in IODP is a closed system, also known as riser drilling. This

445 days 4751 days 507 420 km 1445 holes 5980 m 2111 m

180 880 m 8003 m Latitude 80.47441 N Longitude 8.22731 E Latitude 70.83151 S Longitude 14.57351 W Halifax, Nova Scotia, Canada Pascagoula, Mississippi 9719 tons 2915 tons Seven 16 cyl index Diesel 5@2100 kW (2815 hp) 2@15500 kW (2010 hp) 143 m 21 m 62 m 11 knots 120 days 50 people 65 people 1115 m2 8838 m

technology has been used, in shallow to intermediate water depths, by the offshore oil industry to explore for, and produce oil and gas. Riser drilling uses two pipes: a drill pipe similar to that used for open hole drilling and a wider diameter riser pipe that surrounds the drill pipe and is cemented into the seafloor (Figure 3). The system is closed because drill fluid (seawater and additives) is pumped down the drillpipe (to lubricate the bit and flush rock cuttings away from the bit) and then returned to the ship via the riser. With riser drilling, the drill fluid density can be varied and borehole pressure can be monitored and controlled, thus overcoming the two limitations of open hole drilling. The single limitation of riser drilling is water depth. The current water depth limit of riser technology is approximately 3 km. This water depth limit occurs because the riser, filled with drill fluid and cuttings, puts a large amount of pressure on the rock. This pressure is greater than the strength of the rock; thus, the rock breaks apart under the riser pressure, causing the drilling system to fail.


40


Cuttings returned to ship

Riser Drill pipe

Drilling fluid is pumped down through drill pipe

Drilling fluid and cuttings flow up between the drill pipe and the riser

Seafloor Surface casing

Second casing

Uncased hole

Drilling fluid and cuttings flow up between the drill pipe and the borehole or casing

Drill bit

Figure 3 Diagram of a riser drilling system.

Deep-Sea Sampling Scientific ocean drilling not only requires a borehole, but – more importantly – the recovery of highquality core samples taken as continuously as possible. Recovering sediment and rock samples from below the seafloor in deep water requires the use of wireline tools. Wireline tools are pumped down the center of the drill pipe to the bottom (called the bottom hole assembly) where they are mechanically latched into place near the drill bit in preparation for sampling. Different types of tools are advanced into the geological formation and take a core sample in different ways, depending on the type of sediment or rock. After the tool samples the rock formation, it is unlatched from the bottom hole assembly with a mechanical device, called an overshot, that is sent down the pipe on a wire. The overshot is used to

unlatch the wireline tool, return it to the ship, and recover the core sample. In scientific ocean drilling, three standard wireline core sampling tools are used: the advanced piston corer, the extended core barrel, and the rotary core barrel. The piston corer is advanced into the sediment ahead of the drill bit using pressure applied by shipboard pumps through the drill pipe (Figure 4(a)). The drill fluid pressure in the pipe is increased until the corer shoots into the sediment. After the corer is shot 10 m ahead of the bit, it is recovered using the wireline overshot. The drill pipe and bit are then advanced by rotary drilling another 10 m, in preparation for taking another core sample. The piston corer is designed to recover undisturbed core samples of soft ooze and sediments up to 250 m below the seafloor. In soft to hard sediments, the piston corer can achieve 100% recovery. However, because the cores are taken sequentially, sediment between consecutive cores may not be recovered. To ensure that a continuous sedimentary section is recovered, particularly for paleoclimate studies, the IODP drills a minimum of three boreholes at one site. The positions of the breaks between consecutive core samples are staggered in each borehole so that if sediment is not sampled in one borehole at a core break depth, it will be recovered in the second or third borehole. The extended core barrel is a modification of the oil industry’s rotary corer and is designed to recover core samples of sedimentary rock formation (Figure 4(b)). Typically, the extended core barrel is deployed at depths below the seafloor at which the sediment is too hard for sampling by the piston corer. The extended corer uses the rotation of the drill string to deepen the borehole and cut the core sample. The cutting action is done with a small bit attached to the core barrel. An innovation of this tool is an internal spring that allows the core barrel’s smaller bit to extend ahead of the drill bit in softer formations. In hard formations, where greater cutting action is needed, the spring is compressed and the small core bit rotates with the main drill bit. The rotary core barrel is a direct descendant of the rotary coring system used in the oil industry and is similar to the extended core barrel in its retracted mode. The rotary corer is designed to recover core samples from medium to very hard formations, including igneous rock. The corer uses the rotation of the drill string and the main drill bit to deepen the hole and cut the core sample (Figure 4(c)). In all drilling operations, there are times when the drill pipe must be recovered to the ship, for example, to change a worn bit or to install different bottom hole equipment. Before pulling the pipe out of the



41

Figure 4 Diagrams of the Ocean Drilling Program’s coring tools: (a) advanced piston corer; (b) extended core barrel; and (c) rotary core barrel.

borehole, a re-entry cone (Figure 5) is dropped down the outside of the drill pipe where it free-falls to the seafloor. The cone is used as a guide to re-enter the borehole with a new bit or equipment. The cone is a very small target in deep water (1–6 km) and ancillary tools are needed to locate it for re-entry into the borehole. The cone is first located acoustically, using seafloor transponders. To precisely pinpoint the cone, a video camera is lowered with the drill pipe to visually pinpoint the location and drop the drill pipe back into the borehole. Special corers are used to sample unusual and difficult formations. For example, gas hydrates, ice-like material that is stable under high pressure and low temperature, commonly occur in deep water below the seafloor and require special samplers. Gas hydrates have generated much public interest since they can contain methane gas trapped within their structure, which is thought to be a potential future energy

source. However, when gas hydrate is sampled, it must be kept at in situ pressure conditions to maintain the integrity of the core. Thus, a pressure core sampler is used to sample hydrates. The sampler is similar to the extended corer in that it has its own bit, but it has an internal valve that closes before the sampler is removed from the formation. The closed valve maintains the sample at in situ pressure conditions.

Drilling Measurements Once sampling is completed, logging tools are lowered into the borehole to measure the in situ geophysical and chemical properties of the formation. In IODP, logging tools, developed for use in the oil industry, are most commonly leased from Schlumberger. The origins of logging go back to 1911 when the science of geophysics was new and was just beginning to be used


42


Drill pipe connection Removable line wiper Derrick

Side-entry port for logging tools and wireline Float valve

Moon pool 23′ 4–3/8′′ O.A.L.

Thrusters

Logging tool Drill pipe

5.7 mi

Drill pipe connection Figure 6 Diagram of the side-entry sub technology.

Re-entry cone

Acoustic beacon

Sediment

Hard rock Not to scale

Figure 5 Diagram of the ship, drill pipe, and re-entry cone used in the drilling process.

to explore the internal structure of the Earth. Conrad and Marcel Schlumberger, the founders of Schlumberger, conceived the idea that electrical measurements could be used to detect ore (precious minerals). Working at first alone and then with a number of

associates, they extended the electrical prospecting technique from the surface to the oil well. Now, the use of electric prospecting, called logging, is widely accepted as a standard method in oil exploration. Logging tools are lowered into the borehole using a cable that also transmits the data, in real time, to the ship. The term logging refers to the type of data collected. For example, a borehole log is a record or ledger of the sediment and rock encountered while drilling. Logs are geophysical and chemical records of the borehole. The logging tools typically comprise transmitters and sensors or a sensor alone encased in a robust stainless steel tube. Examples of tool measurements include electrical resistivity, gamma ray attenuation, natural gamma, acoustic velocity, and magnetic susceptibility. Log data provide an almost continuous record of the sediment and rock formation along length of the borehole. Log data are of high quality when collected in the open hole, outside of the drill pipe. The most common method for deploying logging tools is to deploy a device that releases the drill bit from the drill pipe once all drilling and sampling operations are completed. The drill bit falls to the bottom of the borehole and is left there. The drill pipe is retracted and only 75–100 m of pipe is left at the top of the borehole to keep the upper, loose part of the borehole stable. Logging tools are lowered through the drill pipe to the bottom of the borehole. Log data are acquired by slowly raising the tool up the borehole at



a constant speed. When boreholes are unstable and the walls are collapsing into the open hole, another method is used, unique to scientific ocean drilling. In unstable conditions, the drill pipe cannot be retracted to within 75 m of the seafloor without borehole walls collapsing and blocking or bridging the hole. In these situations, after the bit is released, the logging tools are lowered inside the drill pipe to the bottom of the hole. The drill pipe is retracted only enough to expose the logging tools to the open hole, while protecting the remainder of the borehole walls. Then, drill pipe is retracted at the same speed as logging tools are pulled up through the borehole. The technology developed that allows for this unique operation is called the side-entry sub (Figure 6). When inserted as a part of the drill string, the cable, to which the tools are attached, exits the drill pipe at

the side-entry sub, positioned well below the ship. In the way, the logging cable does not interfere with removal of drill pipe. Data in the borehole are also collected using wireline-deployed tools. Sediment temperature is measured with a temperature sensor mounted inside the cutting edge of the piston corer. In addition, other tools are deployed that do not recover a sample. They are pushed into the sediment ahead of the drill bit, left in place for 10–15 min to record temperature, and then pulled back to the ship, where the data are downloaded and analyzed. Special wireline tools have also been used to measure pressure and fluid flow properties of sediment and rock. IODP also uses an oil industry-developed method for logging sediment and igneous rocks – loggingwhile-drilling or LWD. In this method, some of the

Data logger

Re-entry cone

Hydraulic feed through Packer Thermistor string Zone I

43

Screen

Zone II

Optional osmosampler Sinker bar Optional bridge plug screen

Figure 7 Diagram of the advanced CORK system used to seal instruments in the borehole.


44


same wireline logging sensors are repackaged and attached to the drill string, immediately behind the drill bit. The sensors log the geophysical properties of the sediment and rock as drilling proceeds, thus eliminating problems associated with borehole collapse.

Deep Seafloor Observatories The circulation obviation retrofit kit (CORK) is a seafloor observatory that measures pressure, temperature, and fluid composition – important parameters for the study of the dynamics of deep-sea hydrologic systems. CORKs are installed by the IODP for measurements over long periods of time (months to years). Since 1991, observatories have been installed on the deep seafloor in different settings, for example, at mid-ocean ridge hydrothermal systems and at active margins. The CORKs are installed by the drill ship. After a borehole is drilled, a CORK is installed to seal instruments in the borehole away from the overlying ocean (Figure 7). The CORK has two major parts: the CORK body that provides the seal and an instrument cable, for measuring fluid pressures, and temperatures that hangs from the CORK into the borehole. A data recorder is included with the instrument cable. The data recorders have sufficient battery power and memory for up to 5 years of operation. Data are recovered from CORKs using manned submersibles or remotely operated vehicles. The instruments in the CORK measure pressure and temperature spaced along a cable that extends into the sealed borehole. The CORK also includes a valve above the seal where borehole fluids can be sampled. Advanced CORKs are also used to isolate specific and measure the properties of different sediment or rock zones. The IODP installs another type of long-term seafloor observatory for earthquake studies. Seismic monitoring instruments are installed in deep boreholes located in seismically active regions, for example, off the coast of Japan. These data are used to help established predictive measures to prevent loss of life and damage to cities during large earthquakes. Deep-sea seismic observatories contain a strainmeter, two seismometers, a tilmeter, and a temperature sensor. The observatories have replaceable data-recording devices and batteries like CORKs, and are serviced by remotely operated vehicles. Eventually, real-time power supply and data retrieval will be possible when some of the observatories are connected to nearby deep-sea fiber-optic cables.

Summary Deep-sea scientific drilling applies innovative sampling, instrument, and observatory technologies to the study of Earth system science. These range from the study of Earth’s past ocean and climate conditions using high-quality sediment cores, to the study of earthquakes and tectonic processes using logging tools and seafloor observatories, to exploring gas hydrates (a potential future energy source) using specialized sampling tools. In 2004, the IODP succeeded two earlier scientific programs, Deep Sea Drilling and Ocean Drilling. The IODP operates two ships: the D/V JOIDES Resolution and the D/V Chikyu and leases ships for special operations in shallow water and ice-covered seas. The IODP is supported by the US, Japan, and Europe.

See also Deep Submergence, Science of. Deep-Sea Drilling Results. Manned Submersibles, Deep Water. Remotely Operated Vehicles (ROVs).

Further Reading DSDP (1969–86) Initial Reports of the Deep Sea Drilling Project. Washington, DC: US Government Printing Office. http://www.deepseadrilling.org/i_reports.htm (accessed Mar. 2008). Integrated Ocean Drilling Program (2001) Earth, Oceans and Life: Scientific Investigation of the Earth System Using Multiple Drilling Platforms and New Technologies, Integrated Ocean Drilling Program, Initial Science Plan, 2003–2013. Joint Oceanographic Institutions (1996) Understanding Our Dynamic Earth: Ocean Drilling Program Long Range Plan. Washington, DC: Joint Oceanographic Institutions. Oceanography Society (2006) Special Issue: The Impact of the Ocean Drilling Program. Oceanography 19(4). ODP (1985–present) Proceedings of the Ocean Drilling Program, Initial Reports, vols. 101–210. College Station, TX: ODP. http://www-odp.tamu.edu/publications (accessed Mar. 2008). ODP (1985–present) Proceedings of the Ocean Drilling Program, Scientific Results, vols. 101–210. College Station, TX: ODP. http://www-odp.tamu.edu/publications (accessed Mar. 2008). ODP (2004–present) Proceedings of the Integrated Ocean Drilling Program, vols. 300–. College Station, TX: Integrated Ocean Drilling Program. WHOI (1993–94) Oceanus: 25 Years of Ocean Drilling, vol. 36, no. 4. Woods Hole, MA: Woods Hole Oceanographic Institution.


DEEP-SEA DRILLING RESULTS J. G. Baldauf, Texas A&M University, College Station, TX, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 666–676, & 2001, Elsevier Ltd.

Introduction Modern scientific ocean drilling commenced over forty years ago with the inception of Project Mohole. This project was named for its goal of coring a 5– 6 km borehole through thin oceanic crust, continuing through the Mohorovicic Discontinuity and into the earth’s mantle. Project Mohole was active from 1957 through 1966. Although it did not achieve its objective, the Project demonstrated the means for coring in the oceans for scientific purposes and in doing so planted the seed for future decades of scientific ocean drilling.

The current era of scientific coring commenced with the creation of the Joint Oceanographic Institutions for Deep Earth Sampling (JOIDES) in 1964. In 1965, JOIDES completed its first scientific drilling program on the Blake Plateau using the drilling vessel Caldrill. This initial experiment led to the development of the Deep Sea Drilling Project (DSDP) managed by Scripps Institution of Oceanography, California. The scientific objective of DSDP was expanded, compared to that of Project Mohole, and included the recovery of sediments and rocks from throughout the World Ocean to improve the understanding of the natural processes active on this planet. Central to DSDP was the scientific research vessel Glomar Challenger (Figure 1) operated by Global Marine Inc. During the 15 years (1968–1983) of operations, DSDP advanced the scientific frontier by completing 96 scientific expeditions throughout the World Ocean (Figure 2). These expeditions contributed to

Figure 1 Scientific Research Vessel Glomar Challenger operated by Global Marine Inc. for the Deep Sea Drilling Project from 1968 to 1983.


45

46

DEEP-SEA DRILLING RESULTS

advancement of the earth sciences. For example, DSDP confirmed Alfred Wegner’s theory of continental drift, and provided fundamental evidence in support of plate tectonics, established the timing of northern and southern hemisphere glaciations, documented the desiccation of the Mediterranean Sea, documented the northward migration of India and its collision with Asia, determined the history of the major oceanic gateways and their impact on ocean circulation, and improved the understanding of oceanic crust formation and subduction processes. The overwhelming success of DSDP and the numerous scientific questions still remaining provided the framework for the current scientific drilling program, the Ocean Drilling Program (ODP). This new program, established in 1984, continues to explore the history of the earth as recorded in the rocks and sediments beneath the World Ocean. The centerpiece of ODP is the scientific research vessel

JOIDES Resolution (Figure 3) operated by Transocean SedcoForex for Texas AM University. To date 92 scientific expeditions have been completed by ODP (Figure 4).

Science Initiatives The earth system is complex and dynamic with numerous variables and a multitude of forcing and response mechanisms. The sediments preserved beneath the World Ocean record the tempo and variation of the climate system at annual to millennial scales. Likewise, oceanic crust records the environment at the time of its formation. The understanding of these variables and the naturally occurring process active in the Earth’s environment and in the Earth’s interior continue to evolve based on results from scientific ocean drilling. The

80˚N

60˚N

40˚N

20˚N 0˚ 20˚S

40˚S

60˚S

90˚E

120˚

150˚

180˚

210˚

240˚

270˚

300˚

330˚

0˚

30˚

60˚

DSDP Legs 1_ 96, Sites 1_ 624 Figure 2 Geographic location of the sites occupied during scientific expeditions of the Deep Sea Drilling Project from 1968 to 1983.



47

Figure 3 Scientific Research Vessel JOIDES Resolution operated by Transocean SedcoForex for the Ocean Drilling Program 1984 to present.

partitioning of the Earth’s processes into these two themes is in part arbitrary as the external environment is closely linked to that of the Earth’s interior.

Dynamics of the Earth’s Environment The Earth’s climate system consists of processes active in and between five regimes, space, atmosphere, ocean, cryosphere, and crust (Figure 5). For example, variation in incoming solar radiation influences the atmosphere-ocean coupling through changes in precipitation, evaporation and heat exchange. In addition, ice-sheets (and in turn sea level), sea ice, albedo, and terrestrial and oceanic biomass also respond to solar radiation changes. Similarly, variations in the dimensions and shape of the ocean basins, through crust formation on destruction or the opening or closing of oceanic gateways, influence oceanic circulation, as well as the

salinity and temperature of specific water masses. These changes subsequently impact the distribution of nutrients, regional climates, and the Earth’s biomass. The complexity of the climate system is only now beginning to be realized, in part through the ground truthing of climatic models and the historical perspective at various (annual to millennial) scales provide through ocean drilling. ODP’s contributions to understanding the Earth’s environment are extensive. For example, coring off northern Florida provided evidence to support the Cretaceous/Tertiary Boundary meteorite impact theory and its causal impact on widespread extinction of the Earth’s biota. Numerous other cruises have provided insight as to the complexity of the climate system, investigating scientific themes such as the desertification of Africa, implications on climate of the uplift of Tibet plateau, refinement of the glacial history of both hemispheres, including the climate periodicity in icehouse and


48


151

75°N 162

104 105 152 163

60°N

168

146

30°N 138

174A 171B 150 164 172 102 101 109 100 166 174B 165 171A 110 156 170

0°

103 149 173 157

158

128

107

125 190

30°S

131 185

196

106 153

124 108

184

117

154 155 115

116 121

176

199

136 130

192

134

135

194

123 182

118 114

SB

167

143

133

122

179

141

132

146

180

175 112

144

129

195 193

139

191

198

147 111 137 140 148

197

186 126

160

159 142

145

127 161

169

183

189

187

181

177 120

60°S

119 113

178

188

75°S 120°W

90°W

60°W

0°

30°W

30°E

60°E

90°E

150°E

120°E

180°

150°W

Figure 4 Geographic locations of the 91 scientific expeditions currently completed by ODP from 1984 to the present. Legs indicated in black were completed in 1990 and 1991.

SPACE

Space dust

Aerosol atmosphere _ land coupling

CLOUDS

ATMOSPHERE

Precipitation Heat exchange

e.g., CO2 BIOMASS

Changes in solar radiation Terrestrial radiation Evaporation

Wind stress

ODP SEA-ICE

LAND

Changes of atmospheric composition

Atmosphere _ ocean coupling Changes of land features, orography vegetation, albedo, etc.

Atmosphere _ ice coupling

OCEAN

Ice-ocean coupling Sedimentary record is an integrated signal

ICE SHEETS Changes of ocean basin, shape, salinity, temperature, etc.

EARTH

Figure 5 Schematic representation of the major components of the earth’s climatic system. (Adapted from Crowley and North (1992) and the ODP long-range plan (1996).)

greenhouse worlds, and understanding of the history of sea level. In addition, ODP continues to investigate the extent of the biosphere based on evidence of organisms

deep within Earth’s crust. The discovery of organisms in volcanic crust is significant as it extends the depth of the biosphere and indicates that life is possible in extreme environments.



Three scientific themes associated with the dynamics of the Earth’s environment continue to be explored by ODP. These themes include understanding the Earth’s changing climate, the causes and effects of sea level change, and fluids and bacteria as agents of change. Dynamics of the Earth’s Interior

The global cycling of mass and energy in the Earth’s interior and the extrusion of this material into the Earth’s exterior environment impacts global geochemical budgets (Figure 6). For example, the formation of oceanic crust at the mid-ocean spreading centers plays an important role in mantle dynamics, geochemical fluxes, and heat exchange within and between the Earth’s internal and external environments. Similarly, the subduction of a lithospheric plate at a convergent margin results in the melting of the plate, which in turn contributes to the composition and circulation dynamics of the mantle. Associated with subduction zones are volcanic systems through which magma and gases are extruded to Earth’s exterior, contributing to the chemical fluxes of the oceans and atmosphere. Understanding of the processes associated with crustal formation and plate subduction has been enhanced through the recovery of crustal rocks from beneath the oceans. For example, ODP has improved the understanding of the chemical flux by coring within subduction zones to understand the processes

VOLCANIC ARC

associated with subduction, including chemical and mass balances and fluid flow associated with the interface between the two plates, referred to the Decollement zone. Coring of large igneous provinces (LIPs) such as the Kerguelen Plateau and the Ontong Java Plateau has also provided insight into the mantle dynamics and chemical fluxes. These LIPs formed from the injection of magma through volcanic hotspots. Obtaining crustal samples from these regions enhances the understanding of chemical composition and fluxes, mantle dynamics, and the impact of the formation of these features on oceanic and atmospheric chemistry. ODP has also recovered over 500 m of gabbro from a single hole on the South-west Indian Ridge. This sequence represents the most continuous stratigraphic sequence of lower crust from oceanic basement, allowing insight into the chemical composition and structure of oceanic crust. Two other areas of interest for ODP have been the processes associated with the formation of ore bodies and the formation of gas hydrates – frozen methane crystallized with water. Of particular interest is the origin and history of such deposits and their potential influence on global chemical fluxes. Two scientific themes remain central to ODP investigations: exploring the transfer of heat and material to and from earth’s interior and investigating deformation of the lithosphere and earthquake processes.

OCEANIC SEDIMENTS

Formation of ore bodies

49

IGNEOUS OCEAN CRUST

MID-OCEAN SPREADING CENTER Sulfide deposit formed at spreading center

Fate of sediment Assimilation Metasomatism

COLD LITHOSPHERIC PLATE

Melting Melting or dewatering

Metamorphism, dewatering of oceanic crust

Intraplate volcanism

MAGMA Partial melting

High temperature hydrothermal circulation

Fate of subducted ore bodies Metal addition Fate of subducted plate

Figure 6 Schematic representation of the major components of the earth’s tectonic cycle. (Adapted from the ODP long-range plan 1996.)


50


Technological Advances Scientific advances made through ocean drilling are closely linked to advances in technology. New tools such as the development of seafloor observatories, enhanced coring and logging prototypes, seafloor equipment, and state-of-the-art laboratory equipment have provided new opportunities to advance the scientific frontier.

Seafloor Observatories

A reentry cone seal, also referred to as CORK, has been developed and used to seal an ODP hole, thus preventing flow into or out of the borehole. The CORK provides a means for monitoring borehole temperature and pressure as well as recovery of borehole fluid samples. A typical CORK configuration (Figure 7A and B) is characterized by the thermistor string used for the collection of pressure and temperature data, a borehole fluid sampler for collection of borehole fluids, and a data logger for the recording and storage of data from the thermistor string until it can be downloaded via a submersible. The next-generation borehole seals, known as the ACORK, are currently being developed. The ACORK will allow subdivision of the borehole into

isolated segments, allowing monitoring of fluid flow for given horizons or intervals. Coring and Logging Tools

Advanced piston corer (APC) This is a hydraulically actuated piston corer designed to recover undistributed core samples from soft sediments with enhanced core quality and core recovery. Initially developed during the latter phases of DSDP and enhanced during ODP, the APC has become the mainstay for the recovery of highresolution sedimentary records for paleoceanographic and climate studies. Pressure core sampler (PCS) The PCS was developed for the retrieval of core samples from the ocean floor while maintaining near in situ pressures up to 10 000 psi (69 MPa). This tool continues to be critical for investigating gas-bearing sediments, such as gas hydrates, for the analysis of biogeochemical cycling. Formation microscanner (FMS) Adapted from industry, this logging tool provides an oriented, two-dimensional, high-resolution image of the variations in microresistivity around the borehole

Data logger download access Data logger Data logger latch Borehole fluid sampling window

Submersible/ROV platform Reentry cone Seafloor

11¾" casing hanger Cork seals 16" casing hanger 11¾" casing Thermistor string Open borehole (A)

(B)

Figure 7 (A) Schematic of a reentry cone seal, also referred to as a CORK. The tool provides a means for monitoring borehole temperature and pressure as well as recovery of borehole fluid samples. (B) Underwater view of the CORK landed in a reentry cone.



wall. Collected data allow the correlation of coring and logging depth, orientation of cores and location of the cored sections when recovery is less than 100%, mapping of sedimentary structures, and interpretation of depositional environments.

Seafloor Equipment Seafloor equipment such as a reentry cone and the hard rock guide base have been developed and implemented to achieve scientific objectives at specific sites. The reentry cone is a permanent seafloor installation, which serves as a conduit for reentry of the borehole and a platform for supporting various casing strings. Often the drill pipe is pulled from the hole to replace coring bits or to change coring tools. A temporary reentry cone referred to as a ‘free fall funnel’ is also used to allow bit changes. The hard rock guide base (HRGB) was developed to focus the direction of the drill bit into hard, irregular seafloor surfaces that are otherwise undrillable. The difficulty in drilling is both the inability to spud or start a hole as insufficient weight could be applied to the bit and the tendency of the bit to ‘walk’ downhill when trying to start a hole on a sloping surface. The HRGB, when placed on the seafloor, provides support for the drill string to start a hole. Recent technological development of the Hard Rock Reentry System (HRRS) allows a cased reentry hole to be established in bare hard rock without the use of the HRGB.

Laboratory Equipment Mutisensor track (MST)

The MST allows measurements of cores at centimeter scales for examination of changes in physical parameters resulting from changes in lithology, composition, porosity, density, and magnetic susceptibility of the sediment collected. Changes in these properties result from changing oceanographic Table 1 183)

conditions at the time of deposition or from postdeposition processes. The high-resolution records obtained with the MST also allow correlation of data sets from offset holes to ensure completeness of the geological record. In addition, the high-resolution data collected (100–1000 y) can be correlated to data collected through the logging of the borehole to allow direct correlation of laboratory measurements with those from the borehole.

Facilities The scientific research vessel JOIDES Resolution is a dynamically positioned drilling vessel capable of maintaining position over specific locations while coring in water depths down to 8200 m. The vessel was built in Halifax, Nova Scotia, Canada in 1978 and has a length of 143 m, a breadth of 21 m, a gross tonnage of 7539, and a derrick that towers 61.5 m above the water line. A computer-controlled automated dynamic positioning system regulates 12 thrusters in addition to the main propulsion system to maintain the position of the vessel directly over the drill site. JOIDES Resolution was originally operated as an oil exploration vessel. In 1984, ODP converted the vessel into a scientific research vessel by removing the shipboard riser system and adding scientific laboratories. Operating statistics of the JOIDES Resolution are shown in Table 1. Unique to this vessel is the 7-story laboratory structure housing state-of-the-art scientific equipment for use in the studies of geochemistry, microbiology, paleomagnetism, paleontology, petrology, physical properties, sedimentology, downhole measurements, and marine geophysics. This structure also includes support facilities for electronic repair, computers, photography, database management, communications, and conference facilities. Cores and data collected during each scientific cruise are stored at shore-based facilities for future research by members of the international scientific community. ODP maintains four core repositories,

Significant operational highlights of the JOIDES Resolution from 1985 to 1999 (Leg 100–

Deepest hole penetrated Shallowest operational water depth Deepest operational water depth Most core recovered during a single expedition Total number of sites visited Total number of holes cored Total core cored Total core recovered

51

2111 m below the seafloor – Holes 504B completed during Leg 148, eastern Pacific. 37.5 m – Leg 143, north-west Pacific. 5980 m – Leg 129, western Pacific. 8003 m – Leg 175, south-east Atlantic. 535 – ODP Legs 100–183. 1417 – ODP Legs 100–183. 251 017 m – ODP Leg 100 through Leg 183. 170 770 m – ODP Leg 100 through Leg 183.


52


three in North America and one in Europe. The North American repositories include the East Coast Repository (ECR) at Lamont Doherty Earth Observatory (LDEO), which stores cores from the Atlantic Ocean through Leg 150; the Gulf Coast Repository (GCR) at Texas A&M University (TAMU), Texas, which houses cores from the Pacific and Indian Oceans, and the West Coast Repository (WCR) at Scripps Institution of Oceanography, California, which houses cores from the Pacific and Indian Ocean collected during the Deep Sea Drilling Project. The Bremen Core Repository at the University Bremen, Germany, houses cores obtained from the Atlantic since Leg 151. Data collected during each cruise are stored in one of two locations. Downhole measurement (logging) and site survey data are housed at LDEO. All other data collected during a cruise are housed at TAMU.

Scientific Expeditions Using the JOIDES Resolution, ODP recovers sediments or rocks from beneath the World Ocean. Typically two types of coring tools, the Rotary Core Barrel (RCB) and the Advanced Piston Corer (APC), are used to cut the oceanic sediments and rocks. The RCB, typical of the petroleum industry, is used for penetration of hard rocks or sediment, while the APC is used in the less-indurated, softer sediment. The APC tool is typically used to recover the upper several hundred meters of the sediment sequence, with the subsequent sequence cored using the RCB. Once collected, the cores are returned to the ship via a wireline. Once on board the 9.5 m cores are sectioned into 1.5 m lengths for ease of handling within the shipboard laboratories. Following coring operations, the borehole is often logged using standard industrial tools. Logging provides data on the variation in the physical and chemical properties of the sediments and rocks directly from the walls of the borehole. Seafloor laboratories may also be established by instrumenting a borehole with thermistors, water samplers, or seismometers for long term, multiyear monitoring. Each cruise addresses a specific scientific theme based on a rigorous review of proposals submitted by members of the international science community. The duration of a leg depends on the specific objective. Generally each leg is two months in duration. The ship crew consists of about 106 individuals of whom 51 are members of the scientific party. The remaining contingent supports ship and coring operations. The scientific party consists of international

scientists from participating member countries and a technical support staff from TAMU.

International Partnership Eight international members representing 22 countries currently provide funding for the Ocean Drilling Program. The budget for the program is about US$46 million annually with the US National Science Foundation contributing about 66% of the required funds. The remaining funds are provided by five additional full members, each contributing about US$3 million annually, and two associate members each contributing between US$0.5 and 2 million annually. Full members of the program are the Australia/ Canada/Korea/Chinese Taipei Consortium for Ocean Drilling, the European Science Foundation Consortium for Ocean Drilling (Belgium, Denmark, Finland, Iceland, Ireland, Italy, Norway, Portugal, Spain, Sweden, Switzerland, and The Netherlands), Germany, Japan, the United Kingdom, and the United States. Associate members include France and the People’s Republic of China.

Management Structure The ODP management team consists of the Prime Contractor, Joint Oceanographic Institution (JOI), the Science Operator at Texas A&M University, and the Wireline Operator at Lamont Doherty Earth Observatory (Figure 8). As the prime contractor from the National Science Foundation, JOI is responsible for overall management of the program, including scientific planning, operations, and

International funding sources

National Science Foundation ODP Prime Contractor JOI

ODP Science Operator TAMU

Science Advisory Structure _ JOIDES

ODP Wireline Services LDEO

Figure 8 ODP management structure including the National Science Foundation, Joint Oceanographic Institutes (JOI), Texas A&M University (TAMU) and Lamont Doherty Earth Observatory (LDEO). Science advice is provided to JOI through the Joint Oceanographic Institutes for Deep Earth Sampling (JOIDES) panels.



responding to recommendations from the science advisory structure. Texas A&M University is subcontracted through JOI to implement the scientific program developed by the science advisory structure. As Science Operator, TAMU is responsible for shipboard operations, including cruise staffing, maintenance and support of shipboard laboratories, data acquisition, engineering development, publication, core curation, and shipboard logistic (clearance, safety review, etc.) Lamont Doherty Earth Observatory is subcontracted through JOI to implement the wirelinelogging program for each scientific cruise. As Wireline Operator, LDEO is responsible for standard logging, specialized logging, and log analysis support services and database. LDEO is also responsible for the ODP Site Survey Data Bank at LDEO, which archives and distributes site survey data for ODP.

Advisory Structure The ODP advisory structure (Figure 9) provides advice on the scientific program and on logistical activity and program facilities. Scientific guidance is provided by two JOIDES advisory panels, one focused on the Earth’s environment and the second focused on Earth’s interior. The remaining advisory panels provide guidance for shipboard operations and logistics, such as shipboard laboratories, pollution prevention and safety, and site locations. Scientific Guidance

The success of ODP is in the bottom-up approach when it comes to determining the scientific programs JOI EXCOM SCICOM SSEP Environment

SSEP Interior

OPCOM TEDCOM SSP

PPSP

SCIMP

Program Planning & Detailed Planning Groups Figure 9 JOIDES Advisory structure consisting of two science panels (SSEP) and four additional panels technology TEDCOM, Site Survey (SSP), pollution prevention and safety (PPSP) and shipboard measurements (SCIMP) providing guidance and advice to the SCICOM committee (SCICOM) and the Operations subcommittee (OPCOM). Recommendations of SCICOM are forwarded to the Executive committee (EXCOM) for endorsement and to JOI for implementation.

53

most likely to advance the scientific frontier. Individuals or groups of scientists in the international community submit a proposal to answer specific scientific questions. Each proposal is reviewed by the Science Advisory Panels (SSPS) and recommendations are made to the Science Committee (SCICOM), which is responsible for ranking each proposal based on the overall contribution the proposed program would make to understanding Earth history. SCICOM forwards to the Operations Subcommittee (OPCOM), the highest-ranked proposals for scheduling consideration, taking into account available budget, weather constraints, and efficiency of scheduling when comparing days on site versus days in transit between sites. OPCOM sends back to SCICOM a proposed schedule, which is then ratified by SCICOM and forward to the Executive Committee (EXCOM) for approval. Upon approval, EXCOM forwards the schedule to JOI and subsequently to the TAMU and LDEO for implementation. Logistics/Infrastructure

In addition to science planning, the JOIDES Advisory structure provides recommendations on site locations, pollution prevention and safety, shipboard measurements and technological developments. The Site Survey Panel (SSP) is responsible for reviewing all site location data and forwarding recommendations to OPCOM concerning the state of readiness of proposals under consideration for implementation. Similarly, the Pollution Prevention and Safety panel reviews each proposal for safety concerns. This panel is focused on reducing risks associated with the potential occurrence of hydrocarbons to an acceptable level. Both SSP and PPSP make recommendations to OPCOM for consideration when considering proposals for scheduling. The Scientific Measurements Panel (SCIMP) provides advice and guidance on the shipboard laboratories, primarily pertaining to data acquisition, database storage, and data retrieval. In addition, this panel provides recommendations on publications and databases. The Technology Committee (TEDCOM) works closely with the Engineering development teams at TAMU and LDEO to provide guidance on technological developments.

Future Directions Like the Deep Sea Drilling Project before it, the present ODP has a defined duration, with the present program ending 30 September 2003. Although the


54


program as currently configured will end, the need for continued scientific research through ocean drilling remains great. Recognizing this continued demand, the international partners are actively planning for a new program of scientific ocean drilling. The future program is envisioned to utilize multiple platforms, with both Japan and United States providing platforms for the Integrated Ocean Drilling Program (IODP). Japan will provide a new raiser vessel capable of completing deep (43 km) stratigraphic holes. The United States will provide a nonriser vessel to continue work similar to that currently completed by the JOIDES Resolution. In addition to these two platforms, additional alternate platforms will be included for operating in regions not accessible by the other two vessels, such as the Arctic Ocean, shallow water continental margins, and coral reefs, among others. It is envisioned that this new era of scientific ocean drilling will commence in late 2003.

See also Deep Submergence, Science of. Deep-Sea Drilling Methodology. Manned Submersibles, Deep Water. Remotely Operated Vehicles (ROVs).

Further Reading Bleil U and Thiede J (1988) Geological History of the Polar Oceans: Arctic versus Antarctic. NATO Series C, Mathematical and Physical Sciences, vol. 308. Dordrecht: Kluwer Academic.

Coffin MF and Eldholm O (1994) Large igneous provinces: crustal structure, dimensions, and external consequences. Reviews of Geophysics 32: 1--36. Cullen V (1993/94) 25 years of Ocean Drilling. Oceanus 36(4). Dickens GR, Paull CK, Wallace P and the Leg 164 Science party (1997) Direct measurement of in situ methane quantities in a large gas hydrate reservoir. Nature 385: 426–428 Lomask M (1976) Project Mohole and JOIDES, A Minor Miracle, An Informal History of the National Science Foundation, National Science Foundation. Washington, DC: pp. 167–197. Hsu KJ (1992) Challenger at Sea, a Ship that Revolutionized Earth Science. Princeton, NJ: Princeton University Press. Kastner M, Elderfield H, and Martin JB (1991) Fluids in convergent margins: what do we know about their composition, origin, role in diagenesis and importance of oceanic chemical fluxes? Philosophical Transactions of the Royal Society(London) 335: 275--288. Kennett JP and Ingram BL (1995) A 20,000 year record of ocean circulation and climate change from Santa Barbara basin. Nature 377: 510--513. Parkes RJ, Cragg SJ, Bale JM, et al. (1994) Deep bacterial biosphere in Pacific Ocean sediments. Nature 371: 410--413. Summerhayes CP, Prell WL, and Emeis KC (1992) Upwelling Systems: Evolution since the Early Miocene. Geological Society Special Publication, 64. London: Geological Society of London. Summerhayes CP and Shackleton NJ (1986) North Atlantic Paleoceanography. Geological Society Special Publication 21. London: Blackwell Scientific. Warme JE, Douglad RG, and Winterer EL (1981) The Deep Sea Drilling Project: a Decade of Progress, Special Publication 32. Tulsa, OK: Society of Economic Paleontologists and Mineralogists.


DEEP-SEA FAUNA P. V. R. Snelgrove, Memorial University of Newfoundland, St John’s, NL, Canada J. F. Grassle, Rutgers University, New Brunswick, New Jersey, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 676–687, & 2001, Elsevier Ltd.

Overview The deep sea covers more of the Earth’s surface than any other habitat, but because of its remoteness and the difficulty in sampling such great depths, our sampling coverage and understanding of the environment have been limited. There has been a common misperception that the deep sea is species poor, and a commonly used ‘desert’ analogy is hardly surprising given that early sampling found few organisms, and the first deep-sea photographs revealed large plains of rolling hills covered in sediment with little obvious life (Figure 1). Indeed, all lines of evidence suggested that the deep sea is a very inhospitable environment. Temperatures are low (B41C), ambient pressure is extremely high (hundreds of times greater than on land), light is completely absent, and food is generally in very low abundance. But within the last few decades, quantitative samples have revealed what primitive sampling gear and photographs could not – that sediments in the deep sea are teeming with a rich diversity of tiny invertebrates only a few millimeters in size or smaller. These benthic (bottom-dwelling) organisms may reside just above the bottom but closely associated with it (hyperbenthos), on the sediment surface (epifauna), or among the sediment grains (infauna). The change in perception regarding the species richness of the deep sea has continued to evolve; we now know that, on the basis of the combination of species richness per unit area and total size, the deep sea is the most species-rich habitat in the oceans and among the richest on earth. A similar change in perception has occurred with two other generalizations about the deep sea. First, the deep sea is generally thought of as a food-limited environment, where biomass of individuals and communities as a whole are extremely low. Although this generalization usually holds, the surprising discovery of hydrothermal vent communities in the late 1970s, with meter-long tube worms and biomass that rivaled even the most productive shallow-water areas,

proved that clear exceptions exist. A second generalization is that the deep sea has been considered to be an extraordinarily stable habitat, where variables such as salinity, temperature, and food supply are constant, and light and photosynthesis are uniformly absent. Again, this generalization holds in some respects, in that temperature and salinity are often invariant and light is indeed absent. Studies in the last two decades, however, have indicated that smallscale patchiness is common, seasonal variation in phytoplankton production in surface waters can be directly reflected in the material that reaches deepsea sediments, and some deep-sea areas are very dynamic in terms of currents and sediment movement. The recent discoveries in the deep sea raise several interesting questions. First, how can a seemingly inhospitable and physically homogeneous habitat such as the deep sea support a rich diversity of organisms? Second, how can hydrothermal vents support such a high biomass of organisms relative to most deep-sea environments? We now have a firm understanding of the latter question and some definite ideas on the former.

Defining the Habitats Some deep-sea biologists define deep-sea habitats somewhat arbitrarily as those greater than 1000 m in depth. For this review, the deep sea is defined as all benthic habitats beyond the edge of the continental shelf, including the continental slope, continental rise, abyssal plains, ocean ridges (including hydrothermal vents), and deep-ocean trenches. Thus, ocean bottom from B200 to 10 000 m falls within this definition. Most of these regions share the features described above, including low temperature, dependence on organic production 1000s of meters above, high pressure and a sedimentary bottom, but each has unique characteristics as well (Figure 2). The continental slope, because it is adjacent to the continental shelf, generally receives a higher level of organic input than abyssal areas and, with its B31 slope, it is the steepest of the deep-sea environments other than trenches and seamounts and is subsequently subject to occasional sediment slides (called turbidity currents) that may move large volumes of sediment down the slope to the rise and abyssal plains. The continental slope is largely covered in sediments, which often derive from terrestrial and riverine runoff but may also come from marine


55

56

DEEP-SEA FAUNA

Figure 1 Photograph of a typical deep-sea landscape. This photo is from 750 m near St Croix, US Virgin Islands. Infaunal burrows (B), a sea cucumber (C) and a sea whip (W) are visible.

biological production. The continental rise occurs at the base of the slope and can exhibit elevated organic matter relative to lower slope areas because material moving down the slope may accumulate at the rise. By far, the abyssal plains cover the largest portion of the deep sea (B40% of the Earth’s surface), but they also represent the most benign of the deep-sea habitats. Because they are removed from land influence and very deep (h4000 m), the amount of organic matter reaching the bottom is generally quite small, even in comparison with the slope and rise.

Pe gi rcen ve ta n de ge pt at h

Shelf

Open ocean/Deep sea Photic

100 50

0

The deepest ocean habitats are trenches, which form in subduction areas and are characterized by relatively steep sides, poor circulation, and occasional mud slumping. The poor circulation and slumping make trenches particularly inhospitable to most organisms. The sedimentary environment is vertically structured in terms of geochemistry and living organisms. Sediments have limited permeability and oxygen normally penetrates only a few millimeters by diffusion alone. Greater oxygen penetration occurs when bottom currents mix sediments or when organisms move pore water and sediments around (bioturbation). Sediments with active bioturbation are usually oxygenated within the top few centimeters, although strong bioturbation can lead to deeper pockets of penetration. Because light is absent from the deep sea, most productivity is provided by phytoplankton detritus and fecal pellets sinking from surface waters above, or closer to coastal habitats, from organic material transported seaward (e.g. kelps, seagrass etc.). Most of the available organic matter is concentrated near the sediment surface, although some species are capable of ‘caching’ food deeper in the sediment for later use. The combination of limited food and oxygen penetration at depth in sediments results in the vast majority of organisms being confined to the upper few centimeters of

Percentage area above given depth 30

10

50

70

90

1

Sl

4 − 10° C

op

e

< 4° C

Depth (km)

3

Aphotic

Rise 5 Abyssal plain 7

Gravel and exposed bedrock Detrital sand and silt Reducing clay (black)

9

Trench

Oxidizing clay (red) Siliceous ooze Calcareous ooze

Figure 2 Schematic representation of the deep ocean environments and their sedimentary makeup. The horizontal axis has been greatly compressed, and the vertical axis is subsequently exaggerated. (Modified from Wright JE (ed.) (1977). Introduction to the Oceans. Milton Keynes, UK: The Open University.).


DEEP-SEA FAUNA

sediment near the sediment–water interface. Smaller organisms that can tolerate anoxia and larger organisms that maintain a burrow or appendage to the surface can live deeper, but even then distributions are usually only a few centimeters deeper. Historically, the deep sea was perceived to be aseasonal because temperature is largely invariant at deep-sea depths, the absence of any light negates any day-length signal, and the habitat is so far removed from surface waters that it was thought that any signal from surface production would be completely dampened. Evidence in the last two decades has indicated that seasonality is a factor in many deep-sea environments. Samples from a number of different areas around the world and at a full range of depths have shown that the organic content of sediment does change seasonally. The strength of the seasonal signal, not surprisingly, varies with latitude and location; areas with very strong spring blooms are more likely to result in pulses of phytodetritus that sink to the seafloor than areas with weak production cycles. Where pulses are strong, the benthic fauna has been shown to respond quickly to organic input in terms of activity and biomass. Experimental patches of organic enrichment also generate a response by colonizing species. The deep-sea floor lacks the large-scale physical heterogeneity of habitats such as forests and coral reefs, but it is nonetheless far from uniform. Biologically generated features such as burrows and feeding mounds create small-scale heterogeneity that persists for longer periods of time than in shallow water because physical redistribution of sediments by waves does not occur in the deep sea. As organic matter such as phytodetritus sinks to the seafloor and is carried horizontally by currents, small-scale bottom topography creates spatial variation in how that material settles. Depressions on the sea floor, for example, trap phytodetritus. Sessile species such as sea whips, glass sponges and protozoans called xenophyophores co-occur with mobile groups such as sea spiders and sea cucumbers, whose movements across the sediment can create tracks and topography. All of this small-scale heterogeneity acts in concert to create a mosaic of microhabitats for different organisms.

Seamounts Seamounts, like volcanic islands, are mountains that are formed above the ocean floor near spreading centers, and they subsequently break up the landscape of abyssal plains. Because they are generally steep-sided, much of the substrate is volcanic rock,

57

but sedimentary environments occur where sides are not steeply sloped or if the top of the seamount is flattened to form a guyot. Because seamounts can extend large distances above the bottom (thousands of meters), the fauna is often different from that found on the surrounding abyssal plains. The hard substrate, of course, supports very different species from sedimentary environments, and the relatively high flows that can occur over seamounts can support a higher proportion of organisms that feed on suspended particles than most deep-sea environments. In some cases the seamount may extend through the oxygen minimum layer of the Pacific Ocean; these low oxygen conditions favor a very different set of species than areas where oxygen is not limited.

Hydrothermal Vents Hydrothermal vents represent a very specialized and unusual deep-sea environment, and prior to their discovery in 1977, the deep sea was thought to support very low densities of small invertebrates. There had been some debate until that time as to whether high pressure and low temperature constrain the size of benthic organisms, or whether deepsea environments were simply food limited. The discovery of vents quickly answered that question, as researchers discovered dense concentrations of large organisms such as tube worms, clams, and crabs (Figure 3). Compared to the surrounding deep-sea environment, the vents supported extraordinary biomass of large organisms, and led to early analogies of ‘oases in the desert’. From a biomass perspective this is an appropriate analogy, but from a species diversity perspective it is not. Vents occur where tectonic spreading and subduction create fissures in the Earth’s crust, allowing sea water to percolate through the crust and become heated by the mantle (Figure 4). When this water percolates out through the crust again, it is rich in minerals and reduced compounds such as hydrogen sulfide. Water temperature is extremely high (200– 4001C), and is prevented from boiling by the extreme pressure. It cools quickly, however, as it mixes with the ambient sea water that is typically B41C. The mixture of sedimentary and hard substrate habitats that are characteristic of vent fields often support a large biomass of a very specialized fauna. The key to the high productivity of hydrothermal vents are chemoautotrophic bacteria that live freely or form symbioses within specialized tube worms, clams, and mussels. These bacteria utilize hydrogen sulfide to synthesize organic compounds, which in


58

DEEP-SEA FAUNA

A

C

B

Figure 3 Photograph of tube worms (A), clams (B), and polychaete worms (C) from a hydrothermal vent.

Continental crust Trench

Volcanic island

Ocean Oceanic crust

Mantle

Spreading zone (ridge)

Subduction zone Figure 4 Hydrothermal vents occur where plate tectonic spreading and subduction occur. Water percolates through cracks that plate motion creates in the crust and is superheated in the mantle. The high temperatures cause chemical reactions, changing the chemistry of the sea water, and creating a fluid rich in hydrogen sulfide and various compounds.

the latter case may be passed on to the symbiont hosts. Not surprisingly, the hosts are characterized by reduced guts and little or no feeding structures. For both members of the symbiotic partnership, there are clear advantages. The host provides a physically stable habitat in the immediate proximity of hydrogen sulfide and the bacteria provide a rich food supply to the host. But relatively few species can utilize this symbiotic relationship. The extreme temperature gradients and toxic concentrations of hydrogen sulfide create a habitat that few organisms can tolerate. Moreover, the transient nature of vent environments, which generally persist for time scales of only decades, means that organisms must be able to colonize, grow quickly, and reproduce before vent flow ceases.

Defining the Organisms Deep-sea biologists, like other benthic researchers, divide organisms based on size groupings. These groupings are not absolute in that the larval or juvenile stages of one group may be similar in size to adults from a smaller group, but this division is necessary because the sampling logistics that are appropriate for large organisms are inappropriate for small ones. Organisms that can be readily identified in bottom photographs, such as seastars and crabs, are commonly called megafauna (Figure 5A–C). Included here are characteristic deep-sea fish such as grenadiers, which cruise around near the bottom feeding on any falling carcass or disturbing the sediment and creating feeding pits as they feed on


DEEP-SEA FAUNA

A

D

G

B

E

H

C

F

59

Figure 5 Deep-sea organisms, including megafauna (A, rattail fish; B, sea spider; C, brittle star), macrofauna (D, cumacean; E, tanaid; F, polychaete annelid) and meiofauna (G, ostracod; H, nematode). All taxa are from the northwest Atlantic except nematode from San Diego Trough (kindly provided by PJD Lambshead and D Thistle; G Hampson kindly provided photographs D, E, and G).

bottom invertebrates. Some of these taxa migrate up into the water column, providing an additional means by which energy may be cycled between the water column and the benthos. Macrofauna are organisms living on or in the sediments that are retained on a 300-mm sieve; this cutoff is in contrast to the coarser sieves used by shallow-water benthic researchers because the smaller size of deep-sea animals requires a finer sieve. This size grouping includes polychaete annelids, crustaceans, bivalves and many other phyla (Figure 5D–F). Meiofauna are organisms between 40 and 300 mm, and include nematodes, foraminiferans, tiny crustaceans, and many others (Figure 5G–H). Microorganisms pass through a 40-mm sieve and include bacteria and protists. Within any one of these groups, the taxonomic challenges are considerable in that few individuals are trained in taxonomy of deep-sea organisms and many species have yet to be described (see below). Not surprisingly, syntheses across groups are therefore rare in a single study. Feeding modes in deep-sea sediments include omnivores, predators, scavengers and parasites. Most species are deposit feeders that ingest sediment grains and the organic particles and bacteria associated with them; through their feeding activity, these organisms are particularly important bioturbators.

In some areas, suspension feeders filter particles out of the water column above the bottom, but because they rely on suspended particles they are most abundant in energetic environments such as seamounts. As bottom depth increases, both biomass and densities of organisms decrease (Figure 6). Because food resources become scarcer and scarcer as distance from surface water and primary producers increases, this pattern is not unexpected. What is less intuitive, however, is that as food becomes more limiting and densities of organisms decrease, species diversity does not decline.

Sampling the Fauna Early efforts to sample the deep ocean floor used crude trawls towed from surface ships that were ineffective, and undoubtedly contributed to the idea that the environment is species poor. Modern trawls are now used in deep-sea fisheries, and some of these are effective for sampling megafauna living above the sediment or on bedrock. For smaller organisms, Howard Sanders and Robert Hessler, in their important work in the 1960s, used a semiquantitative device called an epibenthic sled comprising a metal frame surrounding a mesh bag. The sled is lowered


60

DEEP-SEA FAUNA

1.0 Meiobenthos

0.5

Wet wt (g m −2 )

0.1

5.0

Macrobenthos

1.0 0.5 0.1

10

2

4

6

Meiobenthos

6

5

Number m −2

10

4

10

Macrobenthos 3

10

2

10

1

10

2

4

6

Depth (km)

Figure 6 Based on Shirayama Y (1983) Size structure of deepsea meio- and macrobenthos in the western Pacific. International Revue der gesamten Hydrobiologie 68: 799–810. Upper panel shows how densities of organisms decline with depth and lower panel shows a similar pattern for biomass.

to the bottom and towed behind a ship, where it skims off the surface sediment and associated fauna (Figure 7C). For hyperbenthos, this is still an instrument of choice, but for infauna it is only semiquantitative and has been replaced by quantitative samplers. In the late 1960s, biologists began using a large metal device called a spade or box corer that is commonly used today to sample macrofaunal organisms. Box corers range in size but typically sample B 0.25 m 0.25 m of ocean bottom. A metal frame surrounds a metal box that is lowered into the sediment from a surface ship (Figure 7A). After the box, which is often subdivided by a grid of metal subcores, enters the sediment, the spade swings down and slices beneath the box, locks in place, and seals the sediment within the corer for the return trip to the surface ship. Although the box corer is effective for sampling macrofauna, it must be handled carefully to avoid a bow wave as it approaches the

sediment. For meiofauna and microbes living right at the sediment surface, even a slight bow wave is a problem because it can blow away the lightest sediments and organisms at the sediment–water interface. To circumvent this problem, a device called a multicorer was developed by the Scottish Marine Biological Association (now the Scottish Association of Marine Sciences). With this sampler, a frame is gently lowered onto the seafloor from a surface ship and individual acrylic cores slowly enter the sediment. Individual cores are typically B 6 cm in diameter and a corer can have 4–12 individual cores or more. Sampling gear that is deployed from surface ships has a distinct disadvantage; although the samples are quantitative, they are largely collected blindly, meaning that there is no frame of reference for the area where the sample was collected. Thus, the corer could land on or near some anomalous feature on the seafloor, and the investigating scientist could have trouble interpreting why the fauna was unusual. The development of research submersibles, such as ALVIN (Figure 7B) or the Johnson Sealink (Figure 7D) allows scientists to actually visit deep-sea environments and watch as their samples are collected at precisely the locations they request. To achieve this, a device called an ALVIN box corer, which is effectively a miniaturized box corer, has been developed. The manipulator arm of the submersible pushes the corer (typically 15 cm 15 cm) into the sediment and then trips the doors that seal-in the sample, much as the spade does for the larger box corer. On board ship, individual subcores are processed over a sieve and then preserved in buffered 4% formaldehyde. Samples are kept in formaldehyde for at least 48 hours and then transferred to 70% ethanol. Hard substrate environments present a different challenge in terms of quantitative sampling. Quantitative removal of hard substrate fauna is near impossible except where the substratum itself may be removed (e.g. manganese nodules). For these environments, visually based surveys, achieved through submersibles, remotely operated vehicles (ROVs), or towed cameras provide the most common means of evaluating fauna. These same approaches are also used for some megafaunal studies in sedimentary environments. The instruments described above are the bases for evaluating faunal abundance, but studies of deep-sea fauna do not always focus on species composition. Physiologists have developed special respiration chambers to study metabolic processes, and ecologists have developed baited traps, colonization trays, and settlement tiles to study species response to


DEEP-SEA FAUNA

A

C

B

D

61

Figure 7 (A) The box corer is one of the basic sampling tools for deep-sea sediments. The corer is lowered from a surface vessel on a wire, and when the box penetrates into the sediment, the spade swings down and slices through the sediment, sealing the sample in the box. Fred Grassle is shown directing the deployment of the corer. (B) ALVIN was one of the first submersibles to be used for deepsea research and played a vital role in the early characterization of hydrothermal vents. (C) An epibenthic sled, developed for deepsea sampling in the 1960s, is towed so that it skims along the seafloor and samples sediments and near-bottom fauna. (D) The Johnson Sealink is a submersible that allows scientists to descend, with a pilot, to the deep-sea floor. ALVIN corers and colonization trays, seen on the front of the submersible, were to be deployed on this dive along with other gear.

resource availability. Most of these instruments are most effectively deployed by submersibles but other approaches have also been used, including free vehicles that are dropped to the bottom from surface ships, and later float to the surface when a release mechanism is triggered.

Patterns of Diversity Although a few scientists before the turn of the century recognized that the deep sea was indeed a species-rich environment, general recognition of this fact did not occur until a series of papers were published by Howard Sanders and Robert Hessler. They collected a series of samples using an epibenthic sled; although this sampling approach is now known to significantly undersample, it represented a marked

improvement over previous gear. Comparison of the data with data from other environments (Figure 8) indicated that the deep sea was among the most diverse of marine sedimentary habitats. From data collected along a transect running from Martha’s Vineyard, Massachusetts to Bermuda, they demonstrated that diversity in deep-sea sediments exceeded that in most shallow areas and rivaled that observed in shallow tropical areas. At the time, this finding represented a startling contrast to current thinking, and even now the desert analogy still persists in some textbooks. More recent work using a box corer (Figure 7A), has shown that the deep sea is not only species rich, but it may also rival tropical rain forests in terms of total species present. Several studies have looked at broad scale pattern in the deep sea, and found that diversity is not


DEEP-SEA FAUNA

uniform with depth, location, or latitude. Michael Rex analyzed patterns with depth in gastropods and other taxa, and found a parabolic pattern, with a peak in diversity between 2000 and 3000 m on the continental slope. Diversity in shallow-water sedimentary habitats, such as in estuaries and tidal flats, is relatively low, then increases somewhat on the continental shelf, peaks along the mid to lower slope and then declines to abyssal plains. A similar pattern has been noted by other researchers, although the slope depth of the diversity peak varies among studies. There have been exceptions noted to this pattern. The abyssal Pacific, for example, has higher diversity than shallower areas, and work in Australia suggests that some coastal environments are extremely diverse, perhaps even exceeding that observed in the deep sea. Studies have examined latitudinal patterns and suggest that diversity in the deep sea may decrease with latitude, at least in the North Atlantic. This work has focused on North Atlantic macrofauna, but other studies on meiofaunal nematodes in the Atlantic and isopod crustaceans from the South Pacific do not support such a trend.

How Many Species Are There? At higher taxonomic levels, the marine environment is inarguably more diverse than any other habitat on Earth, and most of the phyla and classes of organisms that are unique to the marine environment occur in sedimentary environments. Some 90% of all animal families and 28 of 29 nonsymbiont animal phyla occur in marine environments, and of these 29, 13 occur only in marine habitats. Only one animal phylum, the Onychophora, has no living representatives in marine habitats, but even then there are marine fossil forms now known. A number of estimates have been made regarding the possible number of species in the oceans (Figure 9) and considerable debate has arisen from these estimates; much of this controversy arises from the general acceptance that the deep sea has many undescribed species but disagreement over how many. Of particular uncertainty is the validity of assumptions that have gone into various estimates. One approach has been to survey taxonomic specialists who work on different groups, and determine what proportion of their taxon remains undescribed (left panel in Figure 9). The famous tropical rainforest estimate by Terry Erwin (central panel in Figure 9) of 10 million insect species was generated by looking at numbers of beetle species associated with a given species of rainforest tree, estimating what portion of

insects comprises beetles, and then multiplying by estimated numbers of tropical tree species. Based on the rate at which species were added with increased area sampled along a 176-km long depth contour off the eastern United States, Grassle and Maciolek extrapolated to the total area of the deep sea and estimated that there are 10 million deep-sea macrofaunal species. Robert May suggested that because about half of Grassle and Maciolek’s species were previously undescribed, one could extrapolate from the presently described 250 000 marine species to arrive at a projection of B 500 000 total species. Data from the Pacific suggests that only 1 in 20 species has been described; using this estimate, May’s approach would yield B 5 million species. John Lambshead, a nematode ecologist, has noted that meiofaunal nematodes are more abundant and species rich in individual core samples than macrofauna, and his estimate for total nematodes in the deep sea is 100 million species. At present we know little about how widely distributed nematode species may be, making extrapolation even more tenuous than for macrofauna. Species richness comparisons across different environments are complex. Within any given environment, estimating the total numbers of species present is difficult because it is impossible to fully sample the environment. The deep sea is particularly problematic. It has been estimated that of the B 3.25 108 km2 of seafloor that is part of the deep sea, only about 2 km2 has been sampled for macrofauna and 5 m2 has been sampled for meiofauna. Sampling coverage for microbes is poorer still. In addition to the huge area involved, ship time and sample processing are expensive, and relatively few taxonomic specialists know the deep-sea fauna.

Polychaete and bivalve species

62

Tropical shallow Deep sea slope

75

Outer shelf 50 Tropical shallow Tropical estuary

25

Boreal shallow

Boreal estuary 2000 1000 Number of individuals

3000

Figure 8 Comparison of deep-sea communities to other marine sedimentary communities based on plotting numbers of species versus numbers of individuals. (From Sanders HL (1969) Marine benthic diversity and the stability-time hypothesis. Brookhaven Symposium on Biology 22: 71–80.)


DEEP-SEA FAUNA

63

0

Soils and sediments (all taxa)

Tropical insects

Poore and Wilson (1993)

Grassle and Maciolek (1992) May (1992)

Thorson (1971)

Erwin (1982)

3 2

Terrestrial

Terrestrial (described)

Freshwater

1

2

1

10

Freshwater (described)

20

Marine (described) 9 3 10 bacteria? Marine 9 10 nematodes?

30

6

Number of species (× 10 )

Global biodiversity estimates

Deep-sea sediments (macrofauna)

Figure 9 Various estimates of total species richness in marine sediments in comparison with estimates for other environments. Numbered sources in the left panel are: 1Palmer MA, Covich AP, Finlay BJ et al. (1997) Biodiversity and ecosystem processes in freshwater sediments. Ambio 26: 571–577; 2Brussaard L, Behan-Pelletier VM, Bignell DE et al. (1997) Biodiversity and ecosystem functioning in soil. Ambio 26: 563–570; 3Snelgrove PVR, Blackburn TH, Hutchings PA et al. (1997). The importance of marine sedimentary biodiversity in ecosystem processes. Ambio 26: 578–583. For each of these sources, the numbers of described and projected species are given. The arrow in Snelgrove et al. indicates estimates if bacteria (1012?) or nematodes (109?) are included; data on which estimates for these groups are based are very limited. Values in other panels are projected numbers of species. Numbers for marine systems are solid bars. Freshwater and terrestrial refer to species number for all global components of those environments pooled, although estimates for bacteria are not included in these numbers. Additional data sources are Thorson G (1971). Life in the Sea. New York; McGraw-Hill and Erwin TL (1982) Tropical forests: their richness in Coleoptera and other Arthropod species. Coleopterists’ Bulletin 36: 74–75.

Thus, current conclusions on pattern and numbers are based on very limited spatial coverage. Most of the data on species number and pattern has been based on data from the North Atlantic. Clearly the variability in patterns described above suggests that a wider database is needed to effectively test the generality of these concepts. Areas such as the southern oceans that have been poorly sampled in the past offer key pieces to the puzzle of deep-sea pattern. Another shortcoming with large-scale comparisons is that most focus on just one of the size groupings of organisms. Thus, between-group differences in pattern are difficult to attribute to differences in areas sampled or to real differences between the groups. In short, we need deep-sea studies that are broader in geographic coverage and taxonomic coverage before these questions can be resolved definitively.

Low Diversity Environments Not all deep-sea communities are species rich. As described earlier, the hydrogen sulfide and heavy metals emitted at vents are toxic to most species, and species diversity is quite low. Moreover, the short life span of most individual vents makes them among the most unpredictable environments in the deep sea. Like the hard substrate environment around vents, the sediments that occur in hydrothermal vent areas

are inhospitable because of high concentrations of metals, sulfide, and hydrocarbons. There are, nonetheless, often mats of bacteria over these sediments that utilize the hydrogen sulfide as an energy source. Not surprisingly, an increase in hydrothermal flux can quickly ‘cook’ the bacteria, whereas a decrease in flux can starve them to death. A number of other specialized deep-sea environments are species depauperate. Deep-sea trenches are subject to mud slumping and poor circulation as a result of the steep trench sides, and species diversity is very low. Sediments beneath upwelling regions and other highly productive areas, such as the upper slope off Cape Hatteras, are also generally low in diversity because large amounts of organic matter accumulate on the ocean floor and decompose, resulting in hypoxic (low oxygen) conditions that few organisms can tolerate. A few deep-sea areas, such as the ‘HEBBLE’ site on the continental slope off Nova Scotia, are subject to intensive ‘storms’, where currents become intense and sediment resuspension occurs. Evidence suggests that macrofaunal diversity is depressed in such areas, but surprisingly meiofaunal diversity is not. Presumably, the meiofauna are able to cope with the disturbance more effectively than the macrofauna. Although ecological processes are undoubtedly important in the deep sea, evolutionary time scales and processes are also important. Defaunation of


64

DEEP-SEA FAUNA

some deep-sea areas such as the Norwegian Sea over recent geological time scales is thought to have contributed to low diversity. Glaciation defaunated the area by blocking sunlight and reducing circulation, and the shallow sills that surround the basin have likely resulted in very slow reestablishment of deep-sea communities from adjacent basins.

Theories One of the driving questions in deep-sea ecology since the 1960s is how an environment that appears so physically homogeneous is able to support a speciesrich fauna. When Sanders documented the high diversity of deep-sea systems in the late 1960s, he proposed the stability–time hypothesis, in which the high level of stability afforded by the deep sea over evolutionary time has resulted in greater specialization and niche diversification in deep-sea fauna. Certainly the relative stability of deep-sea environments has contributed to the numbers of species present, but if stability were the lone explanation then it is inconsistent with observations of lower diversity on abyssal plains than on adjacent slope habitats. An additional problem is that most species are thought to be relatively nonselective deposit feeders, which is inconsistent with niche specialization. In the early 1970s, several alternative theories were proposed. Predators could prevent competitive equilibria from being attained by infauna by cropping back individuals. But most deep-sea predators appear to be nonselective, and infauna are characterized by slow growth, late reproductive maturity and dominance by older age classes; none of these characteristics would be expected in a predatorcontrolled system. Evidence from shallow-water sediments suggests that at small scales at least, predators decrease sedimentary diversity, but the role that predators may play in maintaining deep-sea diversity remains largely unanswered at this point. The increasing evidence for small-scale spatial and temporal heterogeneity in deep-sea systems has led to speculation that small-scale patches may be important. The patch mosaic model proposes that small patches create disequilibria habitats that promote different species and thus promote coexistence. Studies have tested the patch mosaic model by sampling natural patches or creating experimental patches; both types of study have found that species that occur in most patch types are usually rare or absent from nonpatch sediments. This pattern is consistent with the patch mosaic model, but experiments so far have demonstrated that patches promote only a modest number of species and the overall species

richness in a given patch tends to be lower than in nonpatch areas. It is unclear whether we need to sample more patch types or invoke an altogether different explanation for high diversity. Island biogeography has shown that larger areas tend to support more species, and it has been argued that the large area of the deep sea may be the primary reason for its species richness. To some extent this areal relationship must be a contributing factor, but if area were the only issue then abyssal plains would consistently exceed all other habitats in diversity. Moreover, the deep sea does not appear to add habitat heterogeneity with area as most habitats do; vast areas of sedimentary bottom may sometimes exhibit changes in sediment composition, but even compared to shallow areas the habitat heterogeneity is quite small. Intermediate disturbance has been touted in a number of ecological systems as being important for high diversity. High levels of disturbance can eliminate sensitive species and invariant habitats may allow superior competitors to outcompete weaker species; both scenarios result in reduced diversity. Intermediate disturbance prevents competitive dominants from taking over but is not severe enough to eliminate sensitive species, resulting in high diversity. The strongest evidence for this hypothesis in the deep sea is the mid-slope peak in diversity described earlier, and the reduced diversity observed in disturbed deep-sea habitats such as hydrothermal vents, low oxygen areas, environments with benthic storms, and slumping areas. In summary, we still lack a definitive explanation for which factors are most important in promoting diversity in deep-sea ecosystems. In all likelihood, no single explanation is correct and multiple factors will prove to be important. Efforts are currently underway to try to clarify this question using experimental approaches such as predator exclusion experiments and creation of artificial food patches, along with analytical approaches that analyze pattern with respect to environmental variables. As the available data increase, many of the questions about pattern and cause may become clearer, but there is considerable work to be done.

Threats and Benefits The deep-sea environment has attributes that render it vulnerable to human disturbance, but impacts have nonetheless been modest compared to most marine habitats because of the distance from land, large size, and great depth. Pollutants and nutrients that have created severe problems in coastal areas via land


DEEP-SEA FAUNA

runoff, river runoff or aerosol transport, are usually sufficiently diluted by the time they reach the open ocean that impacts are modest. There is biochemical evidence, however, that pollutants may occur in deep-sea organisms at low concentrations. Fishing, and the habitat destruction it causes, is less widespread in the deep sea than in shallow water because it is expensive and time consuming to fish at great depths. Moreover, the densities of organisms that are present are often insufficient to support commercial fishing. Having said that, there are a number of deepsea fisheries, many of which utilize trawls and dredges that damage the integrity of the benthic habitat, injure organisms, and remove many nontarget species as by-catch. The vulnerability of deepsea species to human activities is well exemplified by fisheries such as that for the Australian orange roughy. Like many deep-sea fisheries, fishing effort has outpaced the capacity of the population to recover, raising the distinct possibility that a sustainable deep-sea fishery may represent an oxymoron. This same vulnerability presumably applies to nontarget species that are removed as by-catch or injured by fishing gear. A third human impact is through waste disposal. Materials ranging from sewage sludge to radioactive waste have been dumped in the deep ocean, largely justified on the basis that the currents are weak so containment is more likely, food chains are far removed from most human harvesting activities, and the large area of the deep sea reduces the likelihood of a major impact. There is also a feeling among some that even if an impact occurs, the organisms that would be affected have little or no obvious economic value to humans. A final potential threat to deep-sea communities is deep-sea mining. There has been some interest in the mining of manganese nodules, small softball-sized nodules that are rich in manganese, nickel and other metals. These nodules occur on the seabed in abyssal plain areas of the oceans, but because of the depths involved, deep-sea mining is not commercially viable at present. Metals such as manganese are, however, of great strategic importance because many countries presently rely on foreign suppliers. Thus, there is a very real chance that deep-sea mining may someday occur. Different mining strategies will, of course, have different types of impact but habitat destruction is likely to be the greatest problem. Because many deep-sea organisms grow very slowly, reproduce at an older age than their shallowwater counterparts, and produce very few offspring per individual, they are thought to be extremely vulnerable to disturbance and habitat damage. In the past it has been assumed, however, that many deepsea species are broadly distributed because there are

65

few barriers to dispersal and little obvious habitat heterogeneity. If this assumption holds, then their vulnerability to extinction might be reduced. At present, our understanding of how quickly species turn over spatially is very limited, particularly for some of the groups like nematodes that have been least studied. With emerging molecular approaches, it is also becoming clear that species that have been treated as cosmopolitan may, in some instances, be species complexes. Given that many deep-sea environments support species that are either of no commercial fishing interest or are not sustainable, is there any reason to exercise caution in how humans impact the deep sea? There are, in fact, several compelling reasons to be concerned. First, the deep sea represents one of the few remaining pristine habitats on Earth. We can say, with only a few exceptions, that deep-sea communities have not been compromised by human development. This attribute makes them one of the last natural laboratories on Earth where the ‘chemicals’ have not been tainted. Because it is so very diverse, the deep sea can provide a natural and uncompromised laboratory in which to test ideas on regulation of biodiversity. A second reason to exercise caution is that the deep sea may represent one of the largest species pools on Earth. From an ethical and esthetic perspective, it could be argued that this characteristic alone is sufficient motivation to limit human disturbance. But from an economic perspective, there is great interest among pharmaceutical companies in organisms with unusual physiologies; the thermophilic bacteria that live at hydrothermal vents, for example, have generated tremendous interest for their bioactive compounds. A third concern is with respect to remediation. Although material dumped in the deep sea may be out of sight and mind, any decision at a later time to remediate (e.g. leaking radioactive waste) would be prohibitively expensive, if it was possible at all. In summary, the deep sea is a vast and relatively undisturbed habitat that may be very vulnerable to human disturbances. Our current understanding of the deep sea and its immense diversity is very limited, but is nonetheless advancing steadily. A precautionary approach will ensure that the unusual attributes of the deep sea, including its rich biodiversity, will not be inadvertently destroyed by ignorance.

See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Demersal Species Fisheries. Macrobenthos. Meiobenthos. Ocean Margin Sediments. Pelagic Biogeography


66

DEEP-SEA FAUNA

Further Reading Gage JD and Tyler PA (1991) Deep-Sea Biology. A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Grassle JF and Maciolek NJ (1992) Deep-sea species richness: regional and local diversity estimates from quantitative bottom samples. American Naturalist 139: 313--341. Gray J, Poore G, and Ugland K (1992) Coastal and deepsea benthic diversities compared. Marine Ecology Progress Series 159: 97--103. Lambshead PJD (1993) Recent developments in marine benthic biodiversity research. Oceanis 19: 5--24. May R (1992) Bottoms up for the oceans. Nature 357: 278--279.

Merrett NR and Haedrich RL (1997) Deep-sea demersal fish and fisheries. London: Chapman Hall. Poore GBC and Wilson GDF (1993) Marine species richness. Nature 361: 597--598. Rex MA (1983) Geographic patterns of species diversity in the deep-sea benthos. In: Rowe GT (ed.) The Sea: DeepSea Biology, vol. 8, pp. 453--472. New York: John Wiley. Rex MA, Stuart CT, and Hessler RR (1993) Global-scale latitudinal patterns of species diversity in the deep-sea benthos. Nature 365: 636--639. Sanders HL and Hessler RR (1969) Ecology of the deep-sea benthos. Science 163: 1419--1424.


DEEP-SEA FISHES J. D. M. Gordon, Scottish Association for Marine Science, Argyll, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 687–693, & 2001, Elsevier Ltd.

Introduction For the purpose of this article a deep-sea fish is one that lives, at least for most of its life, at depths greater than 400 m. The fishes of the continental shelves are usually classified as either pelagic or demersal. These categories are often further subdivided in the deep sea. The pelagic component is comprised of the mesopelagic and bathypelagic fishes that live entirely in the water column and are generally of small adult size. Mesopelagic fishes, e.g. lantern fishes (family Myctophidae) and cyclothonids (family Gonostomatidae), live beneath the photic zone to a depth of approximately 1000 m. Bathypelagic fishes live below 1000 m and are usually highly adapted to life in a food-poor environment. Examples are the deep-water angler fishes (family Ceratidae) and the gulper eels (family Eurypharyngidae). Although the term demersal, referring to fishes living on or close to the bottom, is equally appropriate to the deep sea it has become customary to divide these fish into benthic and benthopelagic species. There is also a trend to refer to these as ‘deep-water’ rather than ‘deep-sea’ fish, thus avoiding the nautical use of deep-sea meaning distance from land. Benthic fishes are those that spend most of their time on the bottom and include the rays (family Rajidae), the flatfishes (e.g. family Pleuronectidae) and the tripod fishes (family Chlorophthalmidae). The benthopelagic fishes are those that swim freely and habitually near the ocean floor and examples include the squalid sharks (family Squalidae), the macrourid fishes (family Macrouridae) and the smoothheads (family Alepocephalidae) (see Figure 1). The long-held belief that deep-sea fish belonged to old (in evolutionary terms) and archaic groups is no longer tenable. The deep sea has been invaded many times and there is little doubt that the specialized pelagic fauna have undergone much of their evolution in the deep sea. On the other hand, the demersal fishes probably evolved mainly in the shallower waters and secondarily invaded the deep sea and therefore, although well-adapted for life at

depth, their special morphological features are usually less well-developed than those found in the meso- and bathy-pelagic fishes. In some areas, such as the deep Norwegian Sea, the colonization by demersal fishes appears to be of fairly recent origin. Although there are very marked regional differences in the deep-water demersal fish faunas there is also a degree of global similarity, and certain families such as the macrourids and smoothheads are dominant. Some species such as several deep-water sharks and the orange roughy (Hoplostethus atlanticus) are widely distributed on continental slopes. Many abyssal species, of which the armed grenadier (Coryphaenoides armatus), the blue hake (Antimora rostrato), and the lizardfishes (Bathysaurus spp.) are good examples, are cosmopolitan in their distribution. Our knowledge of the deep-water demersal fishes ultimately depends on the sampling techniques. The bottom trawl, beam or otter, has been the most widely used sampling method. Each design of trawl is selective for a particular spectrum of fishes, and on the upper continental slope – especially in areas where there is commercial exploitation – a wide range of trawls have been used that has resulted in a better understanding of the fish assemblages. On the lower slope and at abyssal depths fewer, more specialized sampling gears have been utilized, and the assemblages may not be adequately described. Where the seabed is unsuitable for bottom trawling, sampling by longlines and traps has been successful. Advances in submersible technology continue to provide information on the distribution and behavior of deep-water fishes.

Depth-related Changes in Abundance and Biomass The abundance, biomass, and usually the number of species, decreases with increasing depth. This in turn is related to the food supply which, with the exception of a very small contribution of probably o1% from chemosynthesis, is ultimately derived from photosynthesis at the surface, see Figure 2. Surface production is not uniform throughout the world’s oceans, as it depends on the factors necessary for photosynthesis such as light, temperature, and nutrients. In general, it tends to be greatest at higher latitudes where there is strong winter mixing and in areas of upwelling. On the basis of this production,


67

68

DEEP-SEA FISHES

100 cm

Aphanopus carbo

Phycis blennoides

Bathypterois dubius

Hoplostethus atlanticus

Coryphaenoides rupestris

Bathysaurus ferox

Sebastes mentella

Synaphobranchus kaupi

Figure 1 Some deep-water species to show different morphologies.

the upper layer of the world’s oceans can be divided into a series of fairly well-defined faunal provinces, which extend into the mesopelagic. These provinces have been well-defined for some groups such as the

myctophid fishes. Since the demersal fishes ultimately depend on the same energy supply it is not unreasonable to suppose that such faunal provinces exist also for the demersal fishes, but with a few


DEEP-SEA FISHES

69

0m Surface production

Vertical migration

Deep scattering layer

Rain of dead organisms

Continental shelf

Mesopelagic Horizontal and vertical impingement of mesopelagic food

1000 m

Bathypelagic

Overlapping food chains

2000 m

3000 m

Benthopelagic Food fall

4000 m

Figure 2 A diagram showing some of the pathways by which food reaches the deep-sea fish. Vertical migration is probably an important food source for bottom fish at around 1000 m and could account for the increased biomass at this depth.

exceptions the level of sampling has been insufficient to identify their presence. An indication that such faunal provinces may exist and that they might be directly related to surface production comes from some work on the deep-water, abyssal fishes of the north-eastern Atlantic. The investigations were carried out in three main areas, the Porcupine Abyssal Plain (c. 481N; 161W), the Madeiran Abyssal Plain (c. 311N; 251W) and off the west African coast (201N). All the stations were at depths between about 4000 and 5550 m. Striking differences in species composition, morphology, maximum fish size, feeding pattern, and reproductive strategies were demonstrated between the fish catches of these areas. In the area of the Madeiran Abyssal Plain, where there is a well-established seasonal thermocline and surface productivity is relatively low, the abyssal fish fauna is most diverse. The individual fish tend to be of small body size and adapted to life where the food supply is dispersed and limited. The stations off the African coast have a distinct fauna, which is probably attributable to higher productivity caused by upwelling along the continental margin. Further north, where there is a marked seasonal cycle of productivity, the fauna is less diverse and the individual fish have larger body sizes and are adapted to exploiting food sources that tend to be patchily

distributed. There appears to be a general trend for decreasing diversity and increasing body size from low to high latitudes. Overall, the biomass at abyssal depths is low compared with the continental slope and shelf. Another trend that has been reported in the Pacific is for a decrease in fish biomass with increasing distance from the continental land mass. This probably results from increased surface productivity attributable to terrestrial inputs. The abyssal fishes are dependent on the ‘rain’ of detritus and associated bacteria and occasional large food falls for their food supply and as these decrease exponentially from the surface to the seabed, the low biomass is easy to explain. However, at depths of around 1000 m on the continental slopes or around seamounts there is often an increase in demersal fish abundance and biomass. It is this increase in biomass that forms the basis of the developing deep-water fisheries. Most of the fishes at this depth are benthopelagic and studies of their diets (see below) have shown that pelagic and benthopelagic organisms dominate their diet. Many of the prey organisms are vertical migrators, ascending towards the surface at night to feed and descending to a depth of about 1000 m during the day, where they form a deepscattering layer. Where this scattering layer impinges either vertically or horizontally onto the slope, it


70

DEEP-SEA FISHES

provides a rich source of food for the benthopelagic fishes.

The Diet of Deep-water Demersal Fishes In situ investigations in the deep sea are difficult and inevitably this means that much of our knowledge on feeding is derived from stomach content analysis. This has many inherent difficulties, such as the very high percentage of empty or everted stomachs due to the expansion of the gas during recovery from depth in those species that have gas-filled swim bladders. Very often all that remains are hard parts such as vertebrae, squid beaks, and crustacean appendages that become trapped in the lining of the stomach and which can lead to an overestimation of these prey types. The problem of identifying food items, often from fragments, can lead to bias especially when some prey taxa are poorly known, as is often the case in the deep sea. Indeed in some studies, the contents of the fish stomachs have been considered as yet another method of sampling the deep-sea fauna. For example, a significant part of the mesopelagic fauna of Madeira was described from the stomachs of deep-water species such as the black scabbardfish (Aphanopus carbo) landed by the commercial fishery. Net feeding is also considered to be a problem because of the long time the fish spend in the net after initial capture. In some fishes, such as the Alepocephalidae, there is often a high percentage of unidentified soft tissue that may result from feeding on gelatinous plankton. The gelatinous plankton is poorly sampled by nets but the deployment of cameras has shown that it can be abundant and therefore it should not be neglected as a potentially significant food source. Indirect evidence of feeding modes can be obtained from parasite loadings, presence of sediment in the gut, the morphology of the fish, and its associated sensory systems. Stable isotope ratios are beginning to be used to determine the level of the different fishes in the food chain. Direct observation from manned submersibles, remotely operated vehicles, and baited camera systems is a useful tool for understanding feeding behavior. The feeding strategies of the deep-sea fishes cover as wide a range as their shelf counterparts and, as one might expect, reflect the habitat differences (e.g. depth, bottom topography). The piscivorous fishes can broadly be divided into those that adopt a sitand-wait strategy, such as the lizardfish (Bathysaurus) and those that are active predators such as some deep-water sharks and the black scabbardfish. Many deep-water species, including the macrourids,

feed on a mixed diet of the larger pelagic and benthopelagic crustaceans, cephalopods, and small fishes. Others feed on a mixed diet of the smaller benthopelagic organisms and epifauna including amphipods, mysids, and copepods. Again this group can be divided into those that actively forage, such as some of the smaller macrourid fishes, and those that sit and wait, such as the tripod fish (Bathypterois spp.). Feeding directly on the benthos, whether it is browsing at the surface, sifting the sediment for infauna, cropping or even scavenging is relatively unimportant and reflects the relatively low amount of energy that reaches the deep-sea floor.

Sensory Systems Olfaction is well developed in some groups such as the Gadiformes (families Macrouridae and Moridae), some of the sharks, and the synaphobranchid eels. It is probably mostly used for the detection of food, but because it can be sexually dimorphic in some species it may also be used for mate recognition. For example, many male macrourids have larger nostrils than the female. In general the eyes of the benthopelagic fishes do not have the wide range of adaptations to life in the deep sea that are found in the meso- and bathy-pelagic fishes. The eyes tend to remain large and functional and are probably used mainly to detect bioluminescence. Relatively few of the benthopelagic fishes have photophores. Some families have adaptations to maximize the incoming light. In the Alepocephalidae there is a large aphakic space which allows more light to reach the retina from around the edge of the lens, even if it is less focussed. In the squalid sharks and in some teleosts there is a reflective tapetum behind the retina which maximizes the stimulation resulting from the light entering the eye. Some species such as the tripod fish and the forkbeards (Phycis spp.) have developed long fin rays that are sensitive to touch and are used for detecting prey. As may be expected in a dark environment, the lateral line system for detecting movements is particularly well developed in deepwater species and in some species it is particularly elaborate and extends onto the head as a series of canals. The elongate body form of many species is probably an adaptation to increase the sensitivity of the lateral line. In some deep-water species, including many macrourids and gadids of the slope, the males have drumming muscles on their swim bladders for producing sound. This adaptation is absent in the abyssal species.


DEEP-SEA FISHES

71

Buoyancy Fishes in general have evolved many different methods of reducing the energy required to maintain them in the water column and deep-water fishes are no exception. Gas-filled swim bladders are widely used by the benthopelagic fishes of the continental slopes and also in some abyssal fishes, such as the macrourid fishes, where it has been shown that there is a direct relationship between the length of the retia mirabila (the blood supply to the gas gland) and the depth of occurrence. In some species the swim bladder has become filled with low-density lipid, such as wax esters as in the orange roughy. Some gasfilled swim bladders also have considerable amounts of phospholipid and/or cholesterol, although their role in buoyancy control is uncertain. Reduction in body density can be achieved by having lipids distributed throughout the body. The orange roughy has wax esters in a layer beneath the skin and in vacuoles on the head. Density can also be decreased by reducing the ossification of the skeleton and by having a high water content in the tissues, such as in the alepocephalid fishes. The deep-water sharks have very large livers and also generate hydrodynamic lift by their pectoral fins during swimming, as in their shallow-water counterparts.

Longevity The otoliths (earbones) of most deep-water fishes have well defined opaque and transparent zones typical of those found in shallow water where, at least in temperate latitudes, they correspond to seasonal changes in growth and hence can be used to age the fish (Figure 3). In shallow water the broader opaque zones are associated with faster summer growth, but it is not so obvious why fish living in the aseasonal deep sea should have changes in growth rate, unless they are linked to seasonal changes in food availability and/or quality. Although otoliths, and sometimes scales, have been used to estimate age of deepwater species on the assumption that the rings are laid down annually it has seldom been possible to directly validate these age estimates except in juvenile specimens. Radiometric aging, although controversial, has tended to confirm the generally held view that many deep-water species are long-lived. The commercially exploited grenadiers (Macrouridae) can live to about 50 years and the orange roughy in New Zealand waters lives to more than 100 years.

Reproduction The benthopelagic fishes have relatively few of the more extreme reproductive adaptations, such as

Figure 3 An otolith (ear bone) of an abyssal macrourid (Coryphaenoides rupestris) showing growth zones similar to the annual rings in shallow-water fish. If these rings are annual then this fish will be 6 years old.

parasitic males and hermaphroditism that are more frequently found in the meso- and bathy-pelagic fishes. However, hermaphroditism does occur in some groups such as the tripod fishes and live-bearing occurs in the Scorpaenidae (e.g., Sebastes spp.) and in the deep-water squalid sharks. It was a long held belief that the lack of seasonality in the deep sea would result in year-round reproduction. However, it has become increasingly apparent that many of the deep-water species of the continental slopes, especially in areas of the oceans where there are marked seasonal cycles of production at the surface, have well-defined spawning seasons. At greater depths year-round spawning or asynchronous spawning is more common. It is also possible that some abyssal species, such as the armed grenadier, are semelparous (spawn once in a lifetime). Many shallow-water fishes begin spawning before they reach full adult size, however there is some evidence that some deepwater fishes may not mature until they reach adult size, thus partitioning the energy supply first for somatic growth and then for reproduction. There is a lack of information on the egg and larval stages of many of the benthopelagic fish species. For example, the eggs and larvae of the commercially important black scabbardfish are unknown and the eggs of the roundnose grenadier (Coryphaenoides rupestris) another widespread and exploited species in the Atlantic, have only been described from the Skagerrak. On the other hand, the abundance of eggs of the orange roughy in the South Pacific has been used for stock assessments. It is probable that the eggs of most species are pelagic, but although there has been speculation about the eggs rising and hatching in the food-rich waters


72

DEEP-SEA FISHES

associated with the thermocline, there is little evidence to substantiate this. Indeed, recent investigations suggest that the ornamentation of the surface of the eggs of some macrourid species might be an adaptation to restrict the ascent of the eggs through the water column and avoid too wide a dispersal.

See also Bioluminescence. Deep-Sea Fishes. Demersal Species Fisheries. Fish Migration, Vertical. Fish Reproduction. Fish Schooling. Mesopelagic Fishes. Open Ocean Fisheries for Deep-Water Species. Upwelling Ecosystems.

Life Histories The relatively low level of sampling in the deep sea, its restriction to a small number of areas and/or depths, and a general lack of seasonal sampling have all resulted in incomplete life history information. Except where there are special physical features, such as extreme temperature changes with depth (e.g. Norwegian Sea), there is little evidence of zonation in the deep-sea fishes. Instead each species has a depth range which can extend over several thousand meters, as in the cut-throat eel (Synaphobranchus kaupi) or over a few hundred meters, as in the tripod fish (Bathypterois dubius) (both examples from the north-east Atlantic). The ‘bigger-deeper’ phenomenon is a common feature among the deep-water demersal fishes, although it might be more correctly referred to as ‘smaller-shallower’. The juveniles of many of the demersal fishes of the continental slopes live at shallower depths than the adults. While in some regions the horizontal distribution of a species can be well documented there is little information on stock discrimination. With present technology it is difficult to tag, release, and recapture deep-water fishes and therefore there is very little information on the movements of deep-water fishes. Some of the commercial deep-water fisheries exist because they often target spawning aggregations, such as orange roughy in the South Pacific and blue ling (Molva dypterygia) in the North Atlantic. Some of the shark species are often found in single sex shoals, and in the exploited leafscale gulper shark (Centrophorus squamosus) of the North Atlantic the gravid females have never been found. The juveniles of many demersal species have never been found in trawl surveys, which suggests that there are separate nursery grounds or that they occur higher in the water column and are not sampled by bottom trawls.

Further Reading Garter JV Jr, Crabtree RE, and Sulak KJ (1997) Feeding at depth. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 115--193. San Diego: Academic Press. Gordon JDM and Duncan JAR (1985) The ecology of the deep-sea benthic and benthopelagic fish on the slopes of the Rockall Trough, northeastern Atlantic. Progress in Oceanography 15: 37--69. Gordon JDM, Merrett NR, and Haedrich RL (1995) Environmental and biological aspects of slope-dwelling fishes of the North Atlantic Slope. In: Hopper AG (ed.) Deep-water Fisheries of the North Atlantic Oceanic Slope, pp. 1--26. Dordrecht: Kluwer Academic Publishers. Haedrich RL (1997) Distribution and population ecology. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 79--114. San Diego: Academic Press. Marshall NB (1979) Developments in deep-sea biology. Poole: Blandford Press. Mauchline J and Gordon JDM (1991) Oceanic pelagic prey of benthopelagic fish in the benthic boundary layer of a marginal oceanic region. Marine Ecology Progress Series 74: 109--115. Merrett NR (1987) A zone of faunal change in assemblages of abyssal demersal fish in the eastern North Atlantic; a response to seasonality in production? Biological Oceanography 5: 137--151. Merrett NR and Haedrich RL (1997) Deep-sea Demersal Fish and Fisheries. London: Chapman and Hall. Montgomery J and Pankhurst N (1997) Sensory physiology. In: Randall DJ and Farrell AP (eds.) Deepsea Fishes, pp. 325--349. San Diego: Academic Press. Pelster B (1997) Buoyancy at depth. In: Randall DJ and Farrell AP (eds.) Deep-sea Fishes, pp. 195--237. San Diego: Academic Press. Randall DJ and Farrell AP (eds.) (1997) Deep-sea Fishes. San Diego: Academic Press.


DEEP-SEA RIDGES, MICROBIOLOGY A.-L. Reysenbach, Portland State University, Portland, OR, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 693–700, & 2001, Elsevier Ltd.

Introduction Microbes are central to deep-sea ridge ecosystems. Here, in the absence of light energy, the geochemical energy is the primary energy source for microbial growth. These microbes gain energy from oxidation of inorganic compounds such as sulfide and are referred to as chemolithotrophs. Chemolithotrophy or chemosynthesis is the basis of the primary productivity at deep-sea hydrothermal vents, and its discovery challenged our traditional view that all ecosystems were driven by light energy and photosynthesis. The chemolithotrophic microbes are found free-living as well as associated as symbionts with the invertebrates. Additionally, heterotrophic microbes are present that utilize the abundant organic carbon available as a result of the high productivity of these ecosystems. Many of the microbiological discoveries of ocean ridges were gleaned through studying microbial processes or trying to grow novel microbes from these environments. More recently, however, a much more comprehensive picture of the diversity of deep-sea vent microbes is emerging owing to culture-independent diversity assessments. This article will indicate how these different approaches have provided a complementary approach to understanding the dynamics of microbes in deep ocean ridge ecosystems.

Geological Setting Hydrothermal venting occurs both in the terrestrial and marine environments, primarily as a direct result of plate tectonic movement. Divergent boundaries (spreading centers) and convergent margins that produce island arcs are two areas where heat release occurs from the ocean crust, generating high-temperature water activity. At spreading centers, as the tectonic plates are pulled apart, hot molten rock deep within the earth will rise up and fill the gap. Additionally, fissures develop in the crust, and sea water percolates into the crust, reacts with the surrounding rocks, and is heated. This chemically altered sea water will eventually be forced back convectively to the ocean floor as superheated, highly reduced,

hydrothermal fluid rich in gases and minerals (Figure 1) – primary energy sources for microbial growth. Back-arc basins form along active plate margins when old ocean crust is subducted beneath the continental plate. Water can move with the sinking oceanic crust and react with the mantle. As the hydrothermal fluid chemistry is a record of its path within the earth’s crust, these back-arc spreading centers produce heterogeneous hydrothermal fluid chemistry relative to the more stable mid-ocean ridge chemistries. A third basic type of tectonic activity that results in hydrothermal activity comprises ‘hot spots’ and seamounts. Hotspots are places where plumes of hot molten rock may begin deep within the mantle and rise up through the entire mantle and crust. Figure 2 identifies some of the hydrothermal areas that have been studied. Back-arc basins and seamounts are of particular interest for vent biologists as they represent examples of possible island biogeography for vent-related species. The differences in chemistries of these different systems may result in selecting for different physiological types of microbes, but few ecological studies have been addressed to explore such microbial ecological questions in marine vent environments.

Microbial Habitats Deep-sea hydrothermal vents represent one of the most chemically diverse habitats for microbial growth. The geochemical and thermal gradients provide a wide range of possible niches for microbes, with a continuum from oxic to anoxic, pH 3.5 to 8.0, 41C to 4001C, and chemical gradients that mirror the physical gradients (Figure 1). These gradients provide a geochemical disequilibrium, which in turn provides chemical energy for microbial growth. Therefore, there are a multitude of different combinations of electron donors and acceptors and carbon sources for microbial growth. The microbes can be psychrophilic (able to grow best below 151C), mesophilic (growing best between 25 and 401C), or thermophilic (growing best above 451C). Hyperthermophiles are thermophiles that grow best at temperatures above 801C. The mesophilic microbes can be free-living, growing in the cold nutrient-rich sea water surrounding deep-sea vents in areas such as the buoyant plume that results as the hydrothermal fluid rises into the water column and disperses laterally or microbial mats covering sediments and


73

74

DEEP-SEA RIDGES, MICROBIOLOGY

Plume Microbes H2 oxidation → CH4 oxidation → Mn oxidation

Thermophilic chemolithotrophs and heterotrophs Epibionts and symbionts

Mesophilic chemolithotrophs and heterotrophs in microbial mats

HOT Focused WARM Diffuse

Seawater

Addition of H+, Cl –, Fe2+, Mn2+, H4SiO4, 3He, H2S, CH4, CO2, H2, K+, Li+, Pb2+, Ca2+, Cu2+, Zn2+

Subsurface Biosphere

Removal of Mg2+, SO42–

MAGMA 1200 °C

Figure 1 Diagrammatic cross-section of a seafloor spreading center. The vertical is not to scale. The arrows indicate direction of fluid flow.

Southern Explorer Ridge

Lucky Strike Rainbow Broken Spur TAG

Juan de Fuca Ridge Okinawa Trough Gorda Ridge Mariana Trough

d

Logatchev

Atl antic

Galapagos Rift

17°S 18°S 21°S

Central Indian Ridge

Ridge (MAR)

9°N Southern East Pacific Rise 13°S

Bay of Plenty

Carlsberg Ridge

i

Lau Back-Arc Basin

Red Sea

Snake Pit M

Manus

Guaymas Basin 21°N Northern East Pacific Rise 13°N

Menez Gwen

e idg

Chil

nR

e

Ri

t

se

So uth-e

es -w

uth

ast Indian Ridge

So

R ctic Pacific Antar

e idg

Figure 2 World map depicting some of the known marine hydrothermal vent sites.


ia Ind


rocks. Other mesophilic microbes at vents include epibionts that attach to the invertebrates colonizing the vents, or endosymbionts found in specialized intracellular compartments in the invertebrates. Thermophilic microbes are restricted to areas where there is close contact with the high-temperature hydrothermal fluid. This may be in the porous sulfidic rocks or chimney structures, in the mineral sediments covering diffuse flow areas, or perhaps in the subsurface environment. Consider, for example, a cross-section through a porous hydrothermal chimney (Figure 3). From fluid flow and geochemical models one can predict the potential physiological types that can inhabit this thermal environment. However, by far the more abundant niches for microbial growth at deep-sea vents are in the aerobic zone surrounding the high-temperature venting, and it has been estimated that microbial processes other than sulfide oxidation may account for as little as 4% of the total microbial productivity (in terms of energy yield) of the vent ecosystem.

Microbial Diversity at Deep-sea Vents Measures of Diversity

Traditionally, microbial diversity was measured by what could be grown and identified in the laboratory. This approach is heavily biased toward how

Mesophilic heterotrophs H2S oxidizers Methanogens Fe oxidizers Mn oxidizers

successfully we are able to understand the conditions that support growth of deep-sea microbes. However, the use of nucleotide sequence comparisons of the small subunit rRNA (16S rRNA in Archaea and Bacteria), has provided an evolutionary molecule by which all life can be compared. With this emerged the ‘Universal Tree of Life,’ in which all life is placed within a phylogenetic framework of three domains, the Archaea, the Bacteria (both prokaryotic lacking a cell nucleus), and the Eukarya (all eukaryotes). Since the 16S rRNA molecule is present in all Bacteria and Archaea, one can extract DNA from an environmental sample, amplify the gene that encodes for the 16S rRNA using the polymerase chain reaction (PCR), sort the different genes from the different microbes by cloning, and sequence the individual genes that represent the ‘fingerprints’ of each phylogenetically distinct microbe in the environment. These sequences can be placed in the phylogenetic framework of a tree, and the diversity of the microbes in the environment can then be identified without the reliance on trying to grow them in the laboratory. This molecular phylogenetic approach to identifying microbes that have previously resisted cultivation has revolutionized our perception of the microbial world. For example, many symbionts cannot grow without their hosts. All of the endosymbionts of deep-sea hydrothermal invertebrates were identified using these molecular phylogenetic approaches.

Chimney wall

Thermophilic and hyperthermophilic Methanogens ‘Knallgas’ bacteria SO4 reducers Fe oxidizers/reducers Heterotrophs

Sea water

2°C

20_ 50°C

75

100°C

200°C

300°C

Vent fluid

350°C

Figure 3 Schematic representation of a hydrothermal vent chimney showing possible microniches for chemolithotrophic thermophiles. (Modified from McCollom and Shock (1997).)


76


Endosymbionts

The initial discovery of deep-sea hydrothermal vents more than 20 years ago was accompanied by the discovery of a deep-sea oasis for invertebrate life. The invertebrates included large bivalves and giant tubeworms (Riftia pachyptila), and later it was shown that many of the invertebrates also harbored endosymbionts. The high productivity of these deep-sea oases has been linked to the successful associations between the chemoautotrophic endosymbionts and their macroinvertebrate hosts. Most of the endosymbionts are chemolithotrophic sulfur-oxidizing gamma Proteobacteria. Riftia pachyptila, the giant tubeworm, houses its symbionts in a specialized structure called the trophosome. The worm is mouthless and gutless and the densities of the endosymbionts can be up to B3.7 109 cells per gram of trophosome. The endosymbionts require sulfide, oxygen, and carbon dioxide. Sulfide oxidation provides the energy for carbon dioxide fixation. The host supplies the microbes with these gases, whereas the symbiont provides the host with a continuous supply of organic carbon. The host transports high concentrations of oxygen and hydrogen sulfide, in the less toxic HS form, via an unusual hemoglobin. These gases are taken directly via the host’s vascular system to the endosymbionts. Carbon dioxide is transported freely as CO2 or HCO3 in the host’s blood without the aid of hemoglobin. The sulfur-oxidizing endosymbionts associated with many of the deep-sea vent invertebrates highlighted the importance of chemoautotrophy in the overall productivity of hydrothermal ecosystems. In areas where methane is prevalent, another endosymbiotic relationship between a microbe and invertebrate also exists, namely between a giant bathymodiolid mussel and methylotrophic (methaneoxidizing) endosymbionts. The symbionts in this methane-based symbiosis are housed in the gill tissues and the symbionts have the stacked internal membrane structures diagnostic of methylotrophs. In some cases where methane and sulfide are in abundance, such as at the Mid-Atlantic Ridge (Snake Pit), the mussels have two different symbionts in their gills, the sulfur-oxidizing symbionts and the methylotrophs. Clearly this provides the mussels with the metabolic versatility to capitalize on the different potential energy sources that may be available any given time at a vent environment. Using molecular techniques, it has been shown that the symbionts of the bivalves such as the giant clams and mussels are transferred vertically, viz. from adults to offspring via the ovarian tissue and oocytes. In tubeworms, however, there is no evidence

to support the transmission of symbionts from the adult through the eggs and larvae. In this case, it is thought that the tubeworm has to acquire its symbionts from the environment, and the larval tubeworms probably do so by ingestion of free-living forms of the symbionts. So how did the symbiotic relationships evolve in the bivalves? There is evidence from molecular phylogenetic comparisons of host phylogenies and the symbiont phylogenies that in the vesicomyid clam (another symbiotic invertebrate from deep-sea vents) the symbiont and host have coevolved and cospeciated. As additional molecular phylogenetic studies are completed for other host–symbiont relationships, more insight into the population genetics and cospeciation of vent invertebrates will no doubt be gleaned. Epibionts

Several studies have explored the diversity of epibionts associated with hydrothermal invertebrates. Initial microscopic observations of invertebrates, such as with the polychaete Alvinella pompejana and the vent shrimp Rimicaris exoculata, revealed dense colonization of microbiota associated with dorsal integuments (Figure 4) and mouth parts, respectively. In both cases, the identity of these epibionts was determined using culture-independent approaches and 16S rRNA phylogenetic analysis. These epibionts belong to the epsilon Proteobacteria, a group that appears to be prevalent both as epibionts and as

Figure 4 An Alvinella pompejana on a petri plate. Note the white microbial filaments covering the parapodia; these have been identified using 16S rRNA techniques and belong to the epsilon Proteobacteria.



free-living microbes in the hydrothermal vent environment. The role these epibionts have in relation to their hosts is unclear; however, they may have some nutritional benefit for their host. The dominant epibiont associated with Alvinella pompejana showed some interesting variation between the populations colonizing the anterior and posterior parts of the worm. This small variation in 16S rRNA sequence was not enough to designate the two different population types as different species, but the variation may represent a variation in ecological niche. Recently it was proposed that microbiologists take on a ‘natural concept’ for microbial species, in which environmental gradients provide different ecological niches and cause environmentally induced genetic variation resulting in different ecotypes or ecospecies.

77

Figure 5 An extensive microbial mat surrounding a clump of Riftia pachypitila in Guaymas Basin, Mexico. The sampling basket and arm of submersible Alvin are in the foreground. (Photograph by A.-L. Reysenbach, S. Cary, and G.W. Luther.)

Free-living Microbial Diversity

Mesophiles The rising particle- and nutrient-rich hydrothermal fluid forms the hydrothermal plume that can be detected many kilometers away from the actual vent site. This buoyant plume is enriched in reduced compounds and provides ecological niches for a succession of different metabolic types, laterally and further from the site of venting. Initial rapid hydrogen oxidation takes place, then methane oxidation, and finally manganese oxidation. The shifts in the microbial community laterally along the hydrothermal source are being investigated. Immediately surrounding the deep-sea vents is a very diverse and dense (up to 109 cells per ml) community of microbes that range from obligate chemolithotrophs to heterotrophs. Much of the initial research on these communities focused on the sulfur-oxdizers that form visible microbial mats such as those seen at Guaymas Basin (Figure 5). These communities are dominated by large filaments belonging to the genera Beggiatoa and Thiothrix. Like its close relatives Thiomargarita and Thioplaca, Beggiatoa, has very large cells, much of which is devoid of cytoplasm. The vacuolar space is where nitrate can accumulate. These organisms can couple anoxic oxidation of H2S with nitrate reduction, so they can occupy the anoxic–oxic water interface. Several new surprises emerged using a molecular phylogenetic approach to assess the microbial diversity in a mat associated with the active deep-sea hydrothermal system at Pele’s vents (Loihi Seamount, Hawaii). Not only did the epsilon proteobacteria (similar to the epibiont described above) predominate in these mats, but also a wide diversity of Archaea were detected, closely related to Archaea from the marine planktonic environment and coastal

sediments. This study further supported the growing evidence that archaea are ubiquitous and not restricted to extreme environments. Perhaps one of the most significant discoveries made regarding mesophilic microbial activity at deep-sea hydrothermal vents was the discovery of microbes that produce copious amounts of inorganic filaments of sulfur. Arcobacter-related isolates are able to produce filamentous sulfur at the sulfide– oxygen interface of a gradient and are probably responsible for the white flocculent material (floc) that covers fresh basalt within weeks following an eruption. These microbes may also proliferate in the shallow subsurface and produce the sulfur filaments in areas where hydrothermal fluid is mixing with oxygenated sea water. If an eruption occurs, this flocculent material is then dispersed onto the ocean floor (Figure 6). Thermophiles An emerging theme with studies of microbial diversity at deep-sea hydrothermal vents is that the thermophiles that have been obtained in cultures are only a very minor component of the diversity of thermophiles at deep-sea vents. Many of the enrichment cultures have focused on microbes that grow best above 801C (hyperthermophiles). Most of these organisms fall within the Archaea, and include over 20 different reports of new members of the heterotrophic and sulfur-reducing Thermococcales. Other archaea that have been isolated include thermophilic methanogens such as Methanococcus jannaschii (whose genome was recently fully sequenced) and Methanopyrus, and sulfate-reducing archaea of the order Archaeoglobales. Other Archaeal


78


the epsilon Proteobacteria also thrived in this chamber, raising the question whether these prevalent types are perhaps thermophilic inhabitants of deep-sea vents. Additionally, novel archaeal lineages were detected that were related to known iron oxidizers and thermoacidophiles (acid- and heat-loving organisms). In an independent study that explored the archaeal diversity of deep-sea vent chimneys, many novel very deeply diverging lineages were identified. Molecular-based inventories of deep-sea microbial diversity provide the baseline database for rigorous microecological studies at vents. Additionally, these assessments provide guidance for enrichment culturing strategies. Figure 6 Biogenic flocculent material (floc) in the water column surrounding a newly erupted vent field at 91N on the East Pacific rise. (Photograph by R. Haymon and D. Fornari, courtesy T. Shank.)

isolates obtained from deep-sea vents are the facultatively aerobic obligate chemolithoautotroph Pyrolobus fumarii that was obtained from a hydrothermal chimney sample and grows up to 1131C, the highest temperature recorded for a microbe growing in laboratory culture. Surprisingly few thermophilic Bacteria have been obtained from deep-sea vents; however, culture-independent approaches may change this perception. Thermophilic Bacillus and Thermus species have been reported from deep-sea vents, as have members of sheathed thermophilic heterotrophs, the Thermotogales. More recently, and perhaps more significantly, representatives of two new lineages never previously reported from vents have been isolated. One, named Desulfurobacterium, is the only sulfurreducing obligate chemolithotrophic thermophile in the domain Bacteria. Additional isolates of this group have confirmed that it may represent a new order. A second lineage, relatively closely related to the deeply branching Aquificales lineage, has also been isolated and is a microaerophilic hydrogen oxidizer. Interestingly, this Aquificales lineage was first recognized as existing at deep-sea vents by analyzing the diversity associated with the deployment of an in situ growth chamber using a culture-independent approach. This chamber is placed on top of a hydrothermal vent for a predetermined time. The fluid flows through the chamber, and microbes can colonize surfaces that are placed in the chamber. Upon retrieval, the chamber is brought back to the surface and the diversity of organisms that are present in the chamber is analyzed using DNA-based techniques. Several interesting surprises emerged from one such study. Not only were there novel bacterial lineages prevalent in this environment, but

Subsurface Biosphere It has been estimated that the subsurface is the major biosphere for microbes. Deep-sea hydrothermal vents may represent surface manifestations of this biosphere, and offer ‘windows’ into the Earth’s interior ecosystem. The challenges now are to explore the extent of this biosphere. These include ocean drilling of hydrothermal environments, but also monitoring of microbial, geochemical, and biological changes that occur after new eruptions on the ocean floor. Perhaps one of the most spectacular examples of indirect evidence for an extensive subsurface biosphere at deep-sea vents is the initial biogenic sulfur flocs that are seen being emitted from eruptions (Figure 6). As described above, this flocculant material is produced by a mesophilic vibriod microbe that produces strands of filamentous sulfur at the sulfide–oxygen interface.

Deep-sea Hydrothermal Vents and the Origins of Life As the early Earth accreted more than 4.0 billion years ago, it had a hot volcanic environment and a surface bombarded by asteroids. As the Earth started to cool, hydrothermal activity was extensive, and it is estimated that there was three times more heat flow due to hydrothermal activity in the Archaean Earth, than there is today. Some of the first evidence for life dates back to 3.8 Ga and some of the first microfossils date to about 3.5 Ga when there is no evidence for oxygen in the environment. Additionally, the deepest branching lineages within the tree of life are all represented by thermophilic microbes. These microbes may represent modern day analogues of their thermophilic ancestors. If life did originate in a hot hydrothermal environment, it is likely that the rich CO2 environment and geochemical disequilibrium



associated with the hydrothermal venting was an excellent energy source for the evolution of chemolithotrophs. Prior to the evolution of life, it is also possible that some of the first molecules may have evolved in this environment. A patent attorney and chemist, Gunter Wa¨chterhaüser has proposed an elegant theory of how some of life’s precursor molecules could have assembled on positively charged surfaces such as pyrite, an abundant mineral in hydrothermal systems.

Conclusion Deep-sea hydrothermal ecosystems represent a frontier in science. Microbiologists have only begun to explore this new frontier. With the use of molecular techniques, rapid genomic sequencing, and better methods for sampling microbial niches at ridge ecosystems we will gain a much more comprehensive insight into the roles that microbes play in these ecosystems. Furthermore, understanding how these organisms thrive in this hostile deep environment, how they may influence the precipitation of minerals, and how they may become fossilized into the rock has implications in our search for the evidence of life (past or present) on other planets.

See also

79

Vent Fauna, Physiology of. Mid-Ocean Ridge Seismic Structure.

Further Reading Bock GR and Goode JA (eds.) (1996) Evolution of Hydrothermal Ecosystems on Earth (and Mars?). New York: Wiley. Jeanthon C (2000) Molecular ecology of hydrothermal vent microbial communities. Antonie van Leeuwenhoek 77: 117--133. Karl DM (ed.) (1995) The Microbiology of Deep-sea Hydrothermal Vents. Boca Raton, FL: CRC Press. McCollom TM and Shock EL (1997) Geochemical constraints on chemolithoautotrophic metabolism by microorganisms in seafloor hydrothermal systems. Geochimica et Cosmochimices Acta 61: 4375--4391. Van Dover CL (2000) The Ecology of Deep-sea Hydrothermal Vents. Princeton, NJ: Princeton University Press. Wa¨chterhaüser G (1988) Before enzymes and templates. Theory of surface metabolism. Microbiological Reviews 52: 452--484. Ward DM, Ferris MJ, Nold SC, and Bateson MM (1998) A natural view of the microbial biodiversity within hot spring cyanobacterial mat communities. Microbiology and Molecular Biology Reviews 62: 1353--1370. Whitman WB, Coleman DC, and Wiebe WJ (1998) Prokaryotes: the unseen majority. Proceedings of the National Academy of Sciences of the USA 95: 6578--6583.

Hydrothermal Vent Biota. Hydrothermal Vent Deposits. Hydrothermal Vent Ecology. Hydrothermal


DEEP-SEA SEDIMENT DRIFTS D. A. V. Stow, University of Southampton, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 700–709, & 2001, Elsevier Ltd.

variation in deep-sea paleocirculation. As this is closely linked to climate, the drift successions of ocean basins hold one of the best records of past climate change. This clear environmental significance, together with the recognition that sandy contourites are potential reservoirs for deep-sea oil and gas, has spurred much current research in the field.

Introduction The recognition that sediment flux in the deep ocean basins might be influenced by bottom currents driven by thermohaline circulation was first proposed by the German physical oceanographer George Wust in 1936. His, however, was a lone voice, decried by other physical oceanographers and unheard by most geologists. It was not until the 1960s, following pioneering work by the American team of Bruce Heezen and Charlie Hollister, that the concept once more came before a critical scientific community, but this time with combined geological and oceanographic evidence that was irrefutable. A seminal paper of 1966 demonstrated the very significant effects of contour-following bottom currents (also known as contour currents) in shaping sedimentation on the deep continental rise off eastern North America. The deposits of these currents soon became known as contourites, and the very large, elongate sediment bodies made up largely of contourites were termed sediment drifts. Both were the result of semipermanent alongslope processes rather than downslope event processes. The ensuing decade saw a profusion of research on contourites and bottom currents in and beneath the present-day oceans, coupled with their inaccurate identification in ancient rocks exposed on land. By the late 1970s and early 1980s, the present author had helped establish the standard facies models for contourites, and demonstrated the direct link between bottom current strength and nature of the contourite facies, especially grain size. Discrimination was made between contourites and other deep-sea facies, such as turbidites deposited by catastrophic downslope flows and hemipelagites that result from continuous vertical settling in the open ocean. Since then, much progress has been made on the types and distribution of sediment drifts, the nature and variability of bottom currents, and the correct identification of fossil contourites. Of particular importance has been the work at Cambridge University in decoding the often very subtle signatures captured in contourites in terms of

80

Bottom Currents At the present day, deep-ocean bottom water is formed by the cooling and sinking of surface water at high latitudes and the deep slow thermohaline circulation of these polar water masses throughout the world’s ocean (Figure 1 and 2). Antarctic Bottom Water (AABW), the coldest, densest, and hence deepest water in the oceans, forms close to and beneath floating ice shelves around Antarctica, with localized areas of major generation such as the Weddell Sea. Once formed at the surface, partly by cooling and partly as freezing sea water leaves behind water of greater salinity, AABW rapidly descends the continental slope, circulates eastwards around the continent and then flows northwards through deep-ocean gateways into the Pacific, Atlantic and Indian Oceans. Arctic Bottom Water (ABW) forms in the vicinity of the subpolar surface water gyre in the Norwegian and Greenland Seas and then overflows intermittently to the south through narrow gateways across the Scotland–Iceland–Greenland topographic barrier. It mixes with cold deep Labrador Sea water as it flows south along the Greenland–North American continental margin. Above these bottom waters, the ocean basins are compartmentalized into water masses with different temperature, salinity, and density characteristics. Bottom waters generally move very slowly (1–2 cm s1) throughout the ocean basins, but are significantly affected by the Coriolis Force, which results from the Earth’s spin, and by topography. The Coriolis effect is to constrain water masses against the continental slopes on the western margins of basins, where they become restricted and intensified forming distinct Western Boundary Undercurrents that commonly attain velocities of 10–20 cm s1 and exceed 100 cm s1 where the flow is particularly restricted. Topographic flow constriction is greater on steeper slopes as well as through narrow passages or gateways on the deep seafloor. Bottom currents are a semipermanent part of the thermohaline circulation pattern, and sufficiently


DEEP-SEA SEDIMENT DRIFTS

120°E

160°W

80°W

0°

80°E

80°N

80°N

60°N

60°N

40°N

40°N

20°N

20°N

0°

81

0°

20°S

20°S

40°S

40°S

60°S

60°S

120°E

160°E

80°E

0°E

80°E

Figure 1 Global pattern of abyssal circulation. Shaded areas are regions of production of bottom waters. (After Stowet al., 1996).

competent in parts to erode, transport and deposit sediment. They are also highly variable in velocity, direction, and location. Mean flow velocity generally decreases from the core to the margins of the current, where large eddies peel off and move at high angles or in a reverse direction to the main flow. Tidal, seasonal, and less regular periodicities have been recorded during long-term measurements, and complete flow reversals are common. Variation in kinetic energy at the seafloor results in the alternation of short (days to weeks) episodes of high velocity known as benthic storms, and longer periods (weeks to months) of lower velocity. Benthic storms lead to sediment erosion and the resuspension of large volumes of sediment into the bottom nepheloid layer. They appear to correspond to episodes of high surface kinetic energy due to local storms. Deep and intermediate depth water is also formed from relatively warm surface waters that are subject to excessive evaporation at low latitudes, and hence to an increase in relative density. This process is generally most effective in semi-enclosed marginal seas and basins. The Mediterranean Sea is currently the principal source of warm, highly saline, intermediate water, that flows out through the Strait of Gibraltar and then northwards along the Iberian and north European margin. At different periods of Earth history warm saline bottom waters will have been equally or more important than cold water masses.

Contourite Drifts Contourite accumulations can be grouped into five main classes on the basis of their overall

morphology: (I) contourite sheet drifts; (II) elongate mounded drifts; (III) channel-related drifts; (IV) confined drifts; and (V) modified drift–turbidite systems (Table 1, Figure 3). It is important to note, however, that these distinctive morphologies are simply type members within a continuous spectrum, so that all hybrid types may also occur. They are also found at all depths within the oceans, including all deep-water (42000 m) and mid-water (300–2000 m) settings. Those current-controlled sediment bodies that occur in shallower water (50–300 m) on the outer shelf or uppermost slope are not considered contourite drifts sensu stricto. The occurrence and geometry of these different types is controlled principally by five interrelated factors: the morphological context or bathymetric framework; the current velocity and variability; the amount and type of sediment available; the length of time over which the bottom current processes have operated; and modification by interaction with downslope processes and their deposits. Contourite Sheet Drifts

These form extensive very low-relief accumulations, either as part of the fill of basin plains or plastered against the continental margin. They comprise a layer of more or less constant thickness (up to a few hundred meters) that covers a large area, but that demonstrates a very slight decrease in thickness towards its margins, i.e., having a very broad lowmounded geometry. The internal seismofacies is typically one of low amplitude, discontinuous reflectors or, in some parts, is more or less transparent. They may be covered by large fields of sediment waves, as in the case of the South Brazilian and


82


Greenland Iceland Europe

Indian Ocean

Africa

STC AC

AD

P

0 Antarctic

Depth (m)

1000 2000 3000 4000 5000 6000 60˚ 50˚ 40˚ 30˚ 20˚ 10˚ 0˚ 10˚ 20˚ 30˚ 40˚ 50˚ 60˚ 70˚ 80˚ 0

P

STC AC AD

0

1000

1000

2000

2000

3000

3000

4000

4000

5000

5000 6000

6000 40˚ 20˚ 0˚ 60˚ N ABW : Arctic bottom water AABW : Antarctic bottom water AAIW : Antarctic intermediate water AIW : Arctic intermediate water CW : Central water DW : Deep Atlantic water UDW : Upper Atlantic water MDW : Middle deep water LDW : Lower deep water

20˚

40˚

60˚

80˚

S Med : Water from the Mediterranean : Regions of upwelling AC : Antarctic convergence AD : Antarctic divergence STC : Subtropical convergence P : Arctic polar front : Minimum oxygen levels

Figure 2 Bottom water masses in the North Atlantic Ocean (Reproduced from Stow et al., 1996).

Argentinian basins where they are also capped in the central region by giant elongate bifurcated drifts. The different hydrological and morphological contexts define either abyssal sheets or slope sheets (also known as plastered drifts). The former carpet the floors of abyssal plains and other deep-water basins including those of the South Atlantic and the central Rockall trough in the north-east Atlantic. The basin margin relief partially traps the bottom currents and determines a very complex gyratory circulation. Slope sheets occur near the foot of slopes where outwelling or downwelling bottom currents exist, such as in the Gulf of Cadiz as a result of the deep Mediterranean Sea Water outwelling at an intermediate water level into the Atlantic, or around the Antarctic margins as a result of the formation and downwelling of cold AABW. They are also found plastered against the slope at any level, particularly

where gentle relief and smooth topography favors a broad nonfocused bottom current, such as along the Hebrides margin and Scotian margin. Abyssal sheet drifts typically comprise fine-grained contourite facies, including silts and muds, biogenicrich pelagic material, or manganiferous red clay, interbedded with other basin plain facies. Accumulation rates are generally low – around 2–4 cm ky1. Slope sheets are more varied in grain size, composition and rates of accumulation. Thick sandy contourites have been recovered from base-of-slope sheets in the Gulf of Cadiz, and rates of over 20 cm ky1 (1000 years) are found in sandy–muddy contourite sheets on the Hebridean slope. Elongate Mounded Drifts

This type of contourite accumulation is distinctly mounded and elongate in shape with variable



Table 1

Drift morphology, classification and dimensions

Drift type

Subdivisions

Size (km2)

Examples

Contourite sheet drift

Abyssal sheet

105–106

Slope (plastered sheet)

103–104

Argentine basin; Gloria Drift Gulf of Cadiz; Campos margin

Slope (patch) sheets

103

Detached drift

103–105

Eirek drift; Blake drift

Separated drift

103–104

Feni drift; Faro drift

Patch-drift

10–103

North-east Rockall trough

Contourite-fan

103–105

Vema Channel exit Sumba drift; East Chatham rise

Elongated mounded drift

Channelrelated drift

Confined drift

103–105

Modified Extended drift–turbidite turbidite bodies systems

103–104

Sculptured 103–104 turbidite bodies Intercalated Can be very turbidite– extensive contourite bodies

Columbia levee South Brazil Basin; Hikurangi fandrift South-east Weddell Sea Hatteras rise

dimensions: lengths from a few tens of kilometers to over 1000 km, length to width ratios of 2 : 1 to 10 : 1, and thicknesses up to 2 km. They may occur anywhere from the outer shelf/upper slope, such as those east of New Zealand to the abyssal plains, depending on the depth at which the bottom current flows. They are very common throughout the North Atlantic, but also occur in all the other ocean basins and some marginal seas. One or both lateral margins are generally flanked by distinct moats along which the flow axis occurs and which experience intermittent erosion and nondeposition. Elongate drifts associated with channels or confined basins are classified separately. Both the elongation trend and direction of progradation are dependent on an interaction between the local topography, the current system and intensity, and the Coriolis Force. Elongation is generally parallel or subparallel to the margin, with both detached and separated types recognized, but progradation can lead to parts of the drift being elongated almost perpendicular to the margin. Internal seismic character

83

reflects the individual style of progradation, typically with lenticular, convex-upward depositional units overlying a major erosional discontinuity. Fields of migrating sediment waves are common. Sedimentation rates depend very much on the amount and supply of material to the bottom currents. On average, rates are greater than for sheet drifts, being between 2 and 10 cm ky1, but may range from o2 cm ky1 for open ocean pelagic biogenic-rich drifts, to 460 cm ky1, for some marginal drifts (e.g., along the Hebridean margin). The sediment type also varies according to input, including biogenic, volcaniclastic, and terrigenous types. Grain size varies from muddy to sandy as a result of longterm fluctuations in bottom current strength. Channel-related Drifts

This type of contourite deposit is related to deep channels, passageways or gateways through which the bottom circulation is constrained so that flow velocities are markedly increased (e.g., Vema Channel, Kane Gap, Samoan Passage, Almirante Passage, Faroe-Shetland Channel etc.). Gateways are very important narrow conduits that cut across the sills between ocean basins and thereby allow the exchange of deep and intermediate water masses. In addition to significant erosion and scouring of the passage floor, irregular discontinuous sediment bodies are deposited on the floor and flanks of the channel, as axial and lateral patch drifts, and at the downcurrent exit of the channel, as a contourite fan. Patch drifts are typically small (a few tens of square kilometers in area, 10–150 m thick) and either irregular in shape or elongate in the direction of flow. They can be reflector-free or with a more chaotic seismic facies, and may have either a sheet or mounded geometry. Contourite fans are much larger cone-shaped deposits, up to 100 km or more in width and radius and 300 m in thickness (e.g., the Vema contourite fan). Channel floor deposits include patches of coarsegrained (sand and gravel) lag contourites, mud–clast contourites and associated hiatuses that result from substrate erosion, as well as patch drifts of finergrained muddy and silty contourites where current velocities are locally reduced. Manganiferous mud contourites and nodules are also typical in places. Accumulation rates range from very low, due to nondeposition and erosion, to as much as 10 cm ky1 in some patch drifts and contourite fans. Confined Drifts

Relatively few examples are currently known of drifts confined within small basins. These typically


84


TYPE I Contourite sheets Drift sheet

Plastered drift

TYPE II Elongate drifts

Detached drift

Separated drift

TYPE III Channel-related drifts Contourite 'fan'

Lateral and axial patch drifts

TYPE IV Confined drift

TYPE V Modified fan-drift

Figure 3 Contourite drift models. (Modified from Faugeres et al., 1999.)

occur in tectonically active areas, such as the Sumba drift in the Sumba forearc basin of the Indonesian arc system, the Meiji drift in the Aleutian trench and an unnamed drift in the Falkland Trough. Apart from their topographic confinement, the gross seismic character appears similar to mounded elongate drifts with distinct moats along both margins. Sediment type and grain size depend very much on the nature of input to the bottom current system. Modified Drift-turbidite Systems: Process Interaction

The interaction of downslope and alongslope processes and deposits at all scales is the normal

condition on the margins as well as within the central parts of present ocean basins. Interaction with slow pelagic and hemipelagic accumulation is also the norm, but these deposits do not substantially affect the drift type or morphology. Over a relatively long timescale, there has been an alternation of periods during which either downslope or alongslope processes have dominated as a result of variations in climate, sealevel and bottom circulation coupled with basin morphology and margin topography. This has been particularly true since the late Eocene onset of the current period of intense thermohaline circulation, and with the marked alternation of depositional style reflecting glacial–interglacial episodes during the past 2 My (million years). At the scale of the drift deposit, this interaction can have different expressions as exemplified in the following examples. 1. Scotian Margin: regular interbedding of thin muddy contourite sheets deposited during interglacial periods and fine-grained turbidites dominant during glacials; marked asymmetry of channel levees on the Laurentian Fan, with the larger levees and extended tail in the direction of the dominant bottom current flow. 2. Cape Hatteras Margin: complex imbrication of downslope and alongslope deposits on the lower continental rise, that has been referred to as a companion drift-fan. 3. The Chatham–Kermadec Margin: the deep western boundary current in this region scours and erodes the Bounty Fan south of the Chatham Rise and directly incorporates fine-grained material from turbidity currents that have traveled down the Hikurangi Channel. This material, together with hemipelagic material, is swept north from the downstream end of the turbidity current channel to form a fan-drift deposit. 4. West Antarctic Peninsula Margin: eight large sediment mounds, elongated perpendicular to the margin and separated by turbidity current channels, have an asymmetry that indicates construction by entrainment of the suspended load of down-channel turbidity currents within the ambient south-westerly directed bottom currents and their deposition downcurrent. 5. Hebridean Margin: complex pattern of intercalation of downslope (slides, debrites, and turbidites), alongslope contourites and glaciomarine hemipelagites in both time and space; the alongslope distribution of these mixed facies types by the northward-directed slope current has led to the term composite slope-front fan for the Barra Fan.



concentrations of coarser material. They have a siltyclay grain size, poor sorting, and a mixed terrigenous (or volcaniclastic)–biogenic composition. The components are in part local, including a pelagic contribution, and in part far-traveled.

Erosional Discontinuities

The architecture of deposits within a drift is complex, stressing variations of the processes and accumulation rates linked to changes in current activity. In many cases, the history of contourite drift construction is marked by an alternation of periods of sedimentation and erosion or nondeposition, the latter corresponding to a greater instability of and/or a drastic change in current regime. The result is the superposition of depositional units whose general geometry is lenticular and whose limits correspond to major discontinuities, that are more or less strongly erosive. These discontinuities can be traced at the scale of the accumulation as a whole and are marked by a strong-amplitude continuous reflector, commonly marking a change in seisomofacies linked to variation in current strength. Such extensive and synchronous discontinuities are typical of most drifts. The principal characteristics of drifts evident in seismic records are shown in Figure 4.

Silty Contourites

These, which are also referred to as mottled silty contourites commonly show bioturbational mottling to indistinct discontinuous lamination, and are gradationally interbedded with both muddy and sandy contourite facies. Sharp to irregular tops and bases of silty layers are common, together with thin lenses of coarser material. They have a poorly sorted clayey-sandy silt size and a mixed composition. Sandy Contourites

These occur as both thin irregular layers and as much thicker units within the finer-grained facies and are generally thoroughly bioturbated throughout. In some cases, rare primary horizontal and crosslamination is preserved (though partially destroyed by bioturbation), together with irregular erosional contacts and coarser concentrations or lags. The mean grain size is normally no greater than fine sand, and sorting is mostly poor due to bioturbational mixing, but more rarely clean and well-sorted sands occur. Both positive and negative grading may be present. A mixed terrigenous–biogenic composition is typical, with evidence of abrasion, fragmented bioclasts and iron oxide staining.

Contourite Sediment Facies Several different contourite facies can be recognized on the basis of variations in grain size and composition. These are listed and briefly described below and illustrated in Figure 5 and 6.

• • • • • •

Siliciclastic contourites (muddy, silty, sandy and gravel-rich variation) Shale-clast/shale-chip contourites (all compositions possible) Volcaniclastic contourites (muddy–silty–sandy variations) Calcareous biogenic contourites (calcilutite, -siltite, -arenite variations) Siliceous biogenic contourites (mainly sand grade) Manganiferous muddy contourites ( þ manganiferous nodules/pavements)

Gravel-rich Contourites

These are common in drifts at high latitudes as a result of input from ice-rafted material. Under relatively lowvelocity currents, the gravel and coarse sandy material remains as a passive input into the contourite sequence and is not subsequently reworked to any great extent by bottom currents. Gravel lags indicative of more extensive winnowing have been noted from both glacigenic contourites and from shallow straits, narrow moats, and passageways, where gravel pavements are

Muddy Contourites

These are homogeneous, poorly bedded and highly biouturbated, with rare primary lamination (partly destroyed by bioturbation), and irregular winnowed

ED Debris Flow

NQ

TWT 400 ms

P

SW

MRD

NQ

TWT 400 ms

P

SSF SWa

85

North-east Rockall Trough SSF BSD

ED

M MRD

Wyville-Thomson ridge NE

NQ P 10 km

BES 85 / 05-4

Figure 4 Seismic profiles of actual drift systems.


TWT 1000 MS

86


(A) Clastic contourites Muddy bioturbated trace lamination

(F) Biogenic contourites

(B)

(G)

Carbonate sand ± clean, laminated ± bioturbated

Silica sand ± clean, laminated ± bioturbated

Silty-muddy mottled irreg. layers bioturbated (C)

(H) Sandy bioturbated trace lamination

Biogenic mud/silt bioturbated trace lamination

(I) Chemogenic contourites

(D) Micro-brecciated shale-chip layer in muddy contourite

(E) Gravel-lag irregular, poorly-sorted reverse-graded ±muddy ±Fe Mn crust

Fe Mn-muddy contourites ± micronodules ± Fe Mn lamination ± Fe Mn pavements (J) 'Shallow-water' contourites Clastic +/or biogenic laminated and bioturbated. Gradational grainsize variation

Figure 5 Contourite facies models for clastic, biogenic, chemogenic, and ‘shallow-water’ contourites. (Reproduced from Stow et al., 1996.)

locally developed in response to high-velocity bottom current activity Shale-clast or Shale-chip Layers

These have been recognized in both muddy and sandy contourites from relatively few locations. They result from substrate erosion under relatively strong bottom currents, where erosion has led to a firmer substrate and in some cases burrowing on the omission surface has helped to break up the semi-firm muds. Calcareous and Siliceous Biogenic Contourites

These occur in regions of dominant pelagic biogenic input, including open ocean sites and beneath areas of upwelling. In most cases bedding is indistinct, but may be enhanced by cyclic variations in composition, and primary sedimentary structures are poorly developed or absent, in part due to thorough bioturbation as in siliciclastic contourites. In rare cases, the primary lamination appears to have been well preserved. The mean grain size is most commonly silty clay, clayey silt or muddy-sandy, poorly sorted and with a distinct sand-size fraction representing

the coarser biogenic particles that have not been too fragmented during transport. The composition is typically pelagic to hemipelagic, with nannofossils and foraminifera as dominant elements in the calcareous contourites and radiolaria or diatoms dominant in the siliceous facies. Many of the biogenic particles are fragmented and stained with either iron oxides or manganese dioxde. There is a variable admixture of terrigenous or volcaniclastic material. Manganiferous Contourites

These manganiferous or ferromanganiferous-rich horizons are common. This metal enrichment may occur as very fine dispersed particles, as a coating on individual particles of the background sediment, as fine encrusted horizons or laminae, or as micronodules. It has been observed in both muddy and biogenic contourites from several drifts. Bottom-current Influence

It is important to recognize that bottom currents will influence, to a greater or lesser extent, other deepwater sediments, particularly pelagic, hemipelagic,



Figure 6 Photographs of contourite facies from cores drilled through existing drift systems. Vertical scales labelled in cm.


87


turbiditic, and glacigenic, both during and after deposition. Where the influence is marked and deposition occurs in a drift, then the sediment is termed contourite. Where the influence is less severe, so that features of the original deposit type remain dominant, then the sediment is said to have been influenced by bottom currents, as in bottom-current reworked turbidites. Some more-laminated facies, as well as the thin, clean, cross-laminated sands originally described from the north-east American margin, are most likely of this type.

Contourite Sequences and Current Velocity Muddy, silty, and sandy contourites, of siliciclastic, volcaniclastic, or mixed composition, commonly Lithology and structure

occur in composite sequences or partial sequences a few decimeters in thickness. The ideal or complete sequence shows overall negative grading from muddy through silty to sandy contourites and then positive grading back through silty to muddy contourite facies (Figure 7). Such sequences of grain size and facies variation are now widely recognized, although not always fully developed, and are most probably related to long-term fluctuations in the mean current velocity. Not enough data exist to be certain of the timescale of these cycles, though some evidence points towards 5000–20 000 cycles for certain marginal drifts. The occurrence of widespread hiatuses in the deepocean sediment record is best related to episodes of particularly intense bottom currents. More locally,

Mean grain size 4 8 16 32 64 μm

Silt mottles and lenses irregular Horizontal alignment Bioturbated

Mottled silt and mud

Massive irregular sandy pockets Bioturbated Contacts sharp to gradational

Sandy silt

Silt mottles, lenses and irregular layers Bioturbated Contacts variable

Mottled silt and mud 10 (cm)

ease decr city Velo

Cross-laminated bioturbated

Silt

Maximum velocity

Bottom Velocit current y incre ase

Laminated

Positively graded

Bioturbated

Mud

Negatively graded

88

Bioturbated

Mud

0

Bioturbation Discontinuous silt lenses Gradational contact Sharp (irregular) contact Figure 7 Composite contourite facies model showing grain size variation through a mud–silt–sand contourite sequence. (Modified from Stow et al., 1996.)



such strong currents result in significant sediment winnowing and the accumulation of sand, gravel, and shale-clast contourites. Thick units of sandy contourites together with sandy turbidites reworked by the bottom current are potentially important as hydrocarbon reservoirs where suitably buried in association with source rocks. Biogenic contourites typically occur in similar sequences of a decimetric scale that show distinct variation in biogenic/terrigenous ratio, generally linked to the grain size variation. This cyclic facies pattern has a longer timescale, in the few examples from which there is good dating, and is closely analogous to the Milankovitch cyclicity recognized in many pelagic and hemipelagic successions. It is, therefore, believed to be driven by the same mechanism of orbital forcing superimposed on changes in bottom-current velocity. The link between contourite sequences and changes in paleoclimate and paleocirculation is an extremely important one. Where such sequences can be correctly decoded then a more accurate understanding of the paleo-ocean and its environment can be built up.

See also Bottom Water Formation. Nepheloid Layers. Ocean Margin Sediments. Sea Level Change.

89

Further Reading Faugeres JC, Stow DAV, Imbert P, Viana A, and Wynn RB (1999) Seismic features diagnostic of contourite drifts. Marine Geology , 162: 1--38. Heezen BC, Hollister CD, and Ruddiman WF (1966) Shaping the continental rise by deep geostrophic contour currents. Science 152: 502--508. McCave IN, Manighetti B, and Robinson SG (1995) Sortable silt and fine sediment size/composition slicing: parameters for paleocurrent speed and paleoceanography. Paleoceanography 10: 593--610. Nowell ARM and Hollister CD (eds.) (1985) Deep ocean sediment transport – preliminary results of the high energy benthic boundary layer experiment. Marine Geology 66. Pickering KT, Hiscott RN, and Hein FJ (1989) DeepMarine Environments: Clastic Sedimentation and Tectonics. London: Unwin Hyman. Stow DAV and Faugeres JC (eds.) (1993) Contourites and Bottom Currents, Sedimentary Geology, Special Volume 82, 1--310. Stow DAV, Reading HG, Collinson J (1996) Deep seas. In: Reading HG (ed.) Sedimentary Environments and Facies. 3rd edn, pp. 380–442. Blackwell Science Publishers. Stow DAV and Faugeres JC (eds.) (1998) Contourites, turbidites and process interaction. Sedimentary Geology Special Issue 115. Stow DAV and Mayall M (eds.) (2000) Deep-water sedimentary systems: new models for the 21st century. Marine and Petroleum Geology, Special Volume 17.


DEMERSAL SPECIES FISHERIES

This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 718–725, & 2001, Elsevier Ltd.

Introduction Demersal fisheries use a wide variety of fishing methods to catch fish and shellfish on or close to the sea bed. Demersal fisheries are defined by the type of fishing activity, the gear used and the varieties of fish and shellfish which are caught. Catches from demersal fisheries make up a large proportion of the marine harvest used for human consumption and are the most valuable component of fisheries on continental shelves throughout the world. Demersal fisheries have been a major source of human nutrition and commerce for thousands of years. Models of papyrus pair trawlers were found in Egyptian graves dating back 3000 years. The intensity of fishing activity throughout the world, including demersal fisheries, has increased rapidly over the past century, with more fishing vessels, greater engine power, better fishing gear and improved navigational and fish finding aids. Many demersal fisheries are now overexploited and all are in need of careful assessment and management if they are to provide a sustainable harvest. Demersal fisheries are often contrasted with pelagic fisheries, which use different methods to catch fish in midwater and close to the water surface. Demersal species are also contrasted with pelagic species (see relevant sections), but the distinction between them is not always clear. Demersal species frequently occur in mid-water and pelagic species occur close to the seabed, so that ‘demersal’ species are frequently caught in ‘pelagic’ fisheries and ‘pelagic’ species in demersal fisheries. For example Atlantic cod (Gadus morhua), a typical ‘demersal’ species, occurs close to the seabed, but also throughout the water column and in some areas is caught equally in ‘demersal’ and ‘pelagic’ fishing gear. Atlantic herring (Clupea harengus), a typical pelagic species, is frequently caught on the seabed, when it forms large spawning concentrations as it lays its eggs on gravel banks. Total marine production rose steadily from less than 20 million tonnes (Mt) in 1950 to around 100 Mt during the late 1990s (Figure 1). The demersal

90

100 80

Million tonnes

Copyright & 2001 Elsevier Ltd.

fish catch rose from just over 5 Mt in 1950 to around 20 Mt by the early 1970s and has since fluctuated around that level (Figure 2). The proportion of demersal fish in this total has therefore declined over the period 1970–1998. The products of demersal fisheries are mainly used for human consumption. The species caught tend to be relatively large and of high value compared with typical pelagic species, but there are exceptions to such generalizations. For example the industrial (fishmeal) fisheries of the North Sea, which take over

60 40 20

0

50

19

60

90

80

70

19

19

19

19

Year Demersal fish

Pelagic fish

Molluscs

Crustaceans

Other fish Aquatic plants

Figure 1 Total marine landings.

20

Million tonnes

K. Brander, International Council for the Exploration of the Sea (ICES), Copenhagen, Denmark

15 10 5 0

50

19

60

19

80

70

19

19

90

19

Year Other demersal Blue whiting

Argentine hake Atlantic cod

Figure 2 Total demersal fish landings.


Sand eels Alaska pollock


half of the total fish catch, are principally based on low-value demersal species which live in or close to the seabed (sand eels, small gadoids). Demersal fisheries are also often known as groundfish fisheries, but the terms are not exact equivalents because ‘groundfish’ excludes shellfish, which can properly be considered a part of the demersal catch. Shellfish such as shrimps and lobsters constitute the most valuable component of the demersal trawl catch in some areas.

Principal Species Caught in Demersal Fisheries For statistical, population dynamics and fisheries management purposes the catch of each species (or group of species) is recorded separately by FAO (UN Food and Agriculture Organization). The FAO definitions of species, categories and areas are used here. FAO groups the major commercial species into a number of categories, of which, the flatfish (flounders, halibuts, soles) and the shrimps and prawns are entirely demersal. The gadiforms (cods, hakes, haddocks) include some species which are entirely demersal (haddock) and others which are not (blue whiting). The lobsters are demersal, but most species are caught in special fisheries using traps. An exception is the Norway lobster, which is caught in directed trawl fisheries or as a by-catch, as well as being caught in traps.

Table 1

91

The five demersal marine fish with the highest average catches over the decade 1989–98 are Alaska (walleye) pollock, Atlantic cod, sand eels, blue whiting and Argentine hake. These species all spend a considerable proportion of their time in mid-water. Sand eels spend most of their lives on or in the seabed, but might not be regarded as a typical demersal species, being small, relatively short-lived and of low value. Sand eels and eight of the other 20 top ‘species’ in fact consist of more than one biological species, which are not identified separately in the FAO classification (Table 1). The vast majority of demersal fisheries take place on the continental shelves, at depths of less than 200 m. Fisheries for deep-sea species, down to several thousand meters, have only been undertaken for the past few decades, as the technology to do so developed and it became profitable to exploit other species. The species caught in demersal fisheries are often contrasted with pelagic species in textbooks and described as large, long-lived, high-value fish species, with relatively slow growth rates, low variability in recruitment and low mortality. There are so many exceptions to such generalizations that they are likely to be misleading. For example tuna and salmon are large, high-value pelagic species. Three of the main pelagic species in the North Atlantic (herring, mackerel and horse mackerel) are longer lived, have lower mortality rates and lower variability of recruitment than most demersal stocks in that area (Table 2).

Total world catch of 20 top demersal fish species (averaged from 1989–1998)

Common name

Scientific name

Tonnes

Alaska pollock Atlantic cod Sand eels Blue whiting Argentine hake Croakers, drums nei Pacific cod Saithe (Pollock) Sharks, rays, skates, etc. nei Atlantic redfishes nei Norway pout Flatfishes nei Haddock Cape hakes Blue grenadier Sea catfishes nei Atka mackerel Filefishes Patagonian grenadier South Pacific hake Threadfin breams nei

Theragra chalcogramma Gadus morhua Ammodytes spp. Micromesistius poutassou Merluccius hubbsi Sciaenidae Gadus macrocephalus Pollachius virens Elasmobranchii Sebastes spp. Trisopterus esmarkii Pleuronectiformes Melanogrammus aeglefinus Merluccius capensis, M. paradox Macruronus novaezelandiae Ariidae Pleurogrammus azonus Cantherhines ( ¼ Navodon) spp. Macruronus magellanicus Merluccius gayi Nemipterus spp.

3 182 645 2 317 261 1 003 343 628 918 526 573 492 528 425 467 385 227 337 819 318 383 299 145 289 551 273 459 266 854 255 421 242 815 237 843 234 446 230 221 197 911 186 201


92


Table 2 The intensity of fishing is expressed as the average (1988–1997) probability of being caught during the next year. The interannual variability in number of young fish is expressed as the coefficient of variation of recruitment. Species shown are some of the principal demersal and pelagic fish caught in the north-east Atlantic Probability of being Coefficient of variation caught during of recruitment next year Demersal species Cod Haddock Hake Plaice Saithe Sole Whiting

33–64% 19–52% 27% 32–48% 29–42% 28–37% 47–56%

38–65% 70–151% 33% 35–56% 45–56% 15–94% 41–61%

Pelagic species Herring Horse mackerel Mackerel

12–40% 16% 21%

56–63% 40% 41%

Fishing Gears and Fishing Operations A very wide range of fishing gear is used in demersal fisheries (see also Fishing Methods and Fishing Fleets), the main ones being bottom trawls of different kinds, which are dragged along the seabed behind a trawler. Other methods include seine nets, trammel nets, gill nets, set nets, baited lines and longlines, temporary or permanent traps and barriers. Some fisheries and fishermen concentrate exclusively on demersal fishing operations, but many alternate seasonally, or even within a single day’s fishing activity, between different methods. Fishing vessels may be designed specifically for demersal or pelagic fishing or may be multipurpose.

Effects of Demersal Fisheries on the Species They Exploit Most types of demersal fishing operation are nonselective in the sense that they catch a variety of different sizes and species, many of which are of no commercial value and are discarded. Stones, sponges, corals and other epibenthic organisms are frequently caught by bottom trawls and the action of the fishing gear also disturbs the seabed and the benthic community on and within it. Thus in addition to the intended catch, there is unintended disruption or destruction of marine life (see also Ecosystem Effects of Fishing)

The fact that demersal fishing methods are nonselective has important consequences when trying to limit their impact on marine life. There are direct impacts, when organisms are killed or disturbed by fishing, and indirect impacts, when the prey or predators of an organism are removed or its habitat is changed. The resilience or vulnerability of marine organisms to demersal fishing depends on their life history. In areas where intensive demersal fisheries have been operating for decades to centuries the more vulnerable species will have declined a long time ago, often before there were adequate records of their occurrence. For example, demersal fisheries caused a decline in the population of common skate (Raia batis) in the north-east Atlantic and barndoor skate (Raia laevis) in the north-west Atlantic, to the point where they are locally extinct in areas where they were previously common. These large species of elasmobranch have life histories which, in some respects, resemble marine mammals more than they do teleost fish. They do not mature until 11 years old, and lay only a small number of eggs each year. They are vulnerable to most kinds of demersal fishery, including trawls, seines, lines, and shrimp fisheries in shallow water. The selective (evolutionary) pressure exerted by fisheries favors the survival of species which are resilient and abundant. It is difficult to protect species with vulnerable life histories from demersal fisheries and they may be an inevitable casualty of fishing. Some gear modifications, such as separator panels may help and it may be possible to create refuges for vulnerable species through the use of large-scale marine protected areas. Until recently fisheries management ignored such vulnerable species and concentrated on the assessment and management of a few major commercial species. In areas with intensive demersal fisheries the probability that commercial-sized fish will be caught within one year is often greater than 50% and the fisheries therefore have a very great effect on the level and variability in abundance (Table 2). The effect of fishing explains much of the change in abundance of commercial species which has been observed during the few decades for which information is available and the effects of the environment, which are more difficult to estimate, are regarded as introducing ‘noise’, particularly in the survival of young fish. As the length of the observational time series increases and information about the effects of the environment on fish accumulates, it is becoming possible to turn more of the ‘noise’ into signal. It is no longer credible or sensible to ignore environmental effects when evaluating fluctuations in demersal fisheries, but a



Effects of the Environment on Demersal Fisheries

Rate (growth, reproduction)

The term ‘environment’ is used to include all the physical, chemical and biological factors external to the fish, which influence it. Temperature is one of the main environmental factors affecting marine species. Because fish and shellfish are ectotherms, the temperature of the water surrounding them (ambient temperature) governs the rates of their molecular, physiological and behavioral processes. The relationship between temperature and many of these rates processes (growth, reproductive output,

0

15 10 Temperature (˚C)

5

20

25

Landings in thousand tonnes

Figure 3 The relationship between temperature and many rate processes (growth, reproductive output, mortality) is domed. The optimum temperature is species and size specific.

mortality) is domed, with an optimum temperature, which is species and size specific (Figure 3). The effects of variability in temperature are therefore most easily detected at the extremes and apply to populations and fisheries as well as to processes within a single organism. Temperature change may cause particular species to become more or less abundant in the demersal fisheries of an area, without necessarily affecting the aggregate total yield. The cod (Gadus morhua) at Greenland is at the cold limit of its thermal range and provides a good example of the effects of the environment on a demersal fishery; the changes in the fishery for it during the twentieth century are mainly a consequence of changes in temperature (Figure 4). Cod were present only around the southern tip of Greenland until 1917, when a prolonged period of warming resulted in the poleward expansion of the range by about 1000 km during the 1920s and 1930s. Many other boreal marine species also extended their range at the same time and subsequently retreated during the late 1960s, when colder conditions returned. Changes in wind also affect demersal fish in many different ways. Increased wind speed causes mixing of the water column which alters plankton production. The probability of encounter between fish larvae and their prey is altered as turbulence increases. Changes in wind speed and direction affect the transport of water masses and hence of the planktonic stages of fish (eggs and larvae). For example, in some years a large proportion of the fish larvae on Georges Bank are transported into the Mid-Atlantic Bight instead of remaining on the Bank. In some areas, such as the Baltic, the salinity and oxygen levels are very dependent on inflow of oceanic water, which is largely wind driven. Salinity

500

3.0

400

2.5

300

2.0

200

1.5

100

1.0

0 1900

1920

1940

1960

1980

Temperature (˚C)

considerable scientific effort is still needed in order to include such information effectively.

93

0.5 2000

Figure 4 Cod catch and water temperature at West Greenland. Temperature is the running five-year mean of upper layer (0–40 m) values. , local catch 20; —— international catch; – – –, temperature.


94


and oxygen in turn affect the survival of cod eggs and larvae, with major consequences for the biomass of cod in the area. These environmental effects on the early life stages of demersal fish affect their survival and hence the numbers which recruit to the adult population.

World Catches from Demersal Fisheries and the Limits Demersal fisheries occur mainly on the continental shelves (i.e., at depths less than 200 m). This is because shelf areas are much more productive than the open oceans, but also because it is easier to fish at shallower depths, nearer to the coast. The average catch of demersal fish per unit area on northern hemisphere temperate shelves is twice as high as on southern hemisphere temperate shelves and more than five times higher than on tropical shelves (Figure 5). The difference is probably due to nutrient supply. The effects of differences in productive capacity of the biological system on potential yield from demersal fisheries are dealt with elsewhere (see also Ecosystem Effects of Fishing). Demersal fisheries provide the bulk of fish and shellfish for direct human consumption. The steady increase in the world catch of demersal fish species ended in the early 1970s and has fluctuated around 20 Mt since then (Figure 2). Many of the fisheries are overexploited and yields from them are declining. In a few cases it would seem that the decline has been arrested and the goal of managing for a sustainable harvest may be closer. 2.5

A recent analysis classified the top 200 marine fish species, accounting for 77% of world marine fish production, into four groups – undeveloped, developing, mature and declining (senescent). The proportional change in these groups over the second half of the twentieth century (Figure 6) shows how fishing has intensified, so that by 1994 35% of the fish stocks were in the declining phase, compared with 25% mature and 40% developing. Other analyses reach similar conclusions – that roughly twothirds of marine fish stocks are fully exploited or overexploited and that effective management is needed to stabilize current catch levels. Fish farming (aquaculture) is regarded as one of the principal means of increasing world fish production, but one should recall that a considerable proportion of the diet of farmed fish is supplied by demersal fisheries on species such as sand eel and Norway pout. Fishmeal is also used to feed terrestrial farmed animals.

Management of Demersal Fisheries The purposes of managing demersal fisheries can be categorized as biological, economic and social. Biological goals used to be set in terms of maximum sustainable yield of a few main species, but a broader and more cautious approach is now being introduced, which includes consideration of the ecosystem within which these species are produced and which takes account of the uncertainty in our assessment of the consequences of our activities. The formulation of biological goals is evolving, but even the most basic, such as avoiding extinction of species, are not being

Northern temperate

2.0 Southern temperate

t/km2

1.5

1.0

Tropical (and Mediterranean)

0.5

0.0 E E W W l. N tl. N c. N c. N A Pa Pa

At

W SE W l. S ac. c. S a P P

At

S C C C EC .E W IO IO. . + B c. E c. W l. W Atl. a a d P P e M

At

Figure 5 Demersal fish landings per unit area of continental shelf o200 m deep for the main temperate and tropical areas. The boxes show the spread between the upper and lower quartiles of annual landings and the whiskers show the highest and lowest annual landings 1950–1998. Atl, Atlantic; Pac, Pacific; IO, Indian Ocean; Med, Mediterranean; BS, Black Sea.



95

100 Phase IV _ Senescent

90

Percentage of resources

80 70

Phase III _ Mature

60 50 40

Phase II _ Developing

30 20

Phase I _ Undeveloped

1991

1993

1989

1987

1985

1981

1983

1977

1979

1975

1971

1973

1969

1967

1963

1965

1961

1957

1959

1953

1955

0

1951

10

Figure 6 Temporal change in the level of exploitation of the top 200 marine fish species, showing the progression from mainly undeveloped or developing in 1951 to mainly fully exploited or overexploited in 1994.

achieved in many cases. At a global level it is evident that economic goals are not being achieved, because the capital and operating costs of marine fisheries are about 1.8 times higher than the gross revenue. There are innumerable examples of adverse social impacts of changes in fisheries, often caused by the effects of larger, industrial fishing operations on the quality of life and standard of living of small-scale fishing communities. Clearly there is scope for improvement in fisheries management. From this rather pessimistic analysis of where fisheries management has got us to date, it follows that a description of existing management regimes is a record of current practice rather than a record of successful practice. Biological management of demersal fisheries has developed mainly from a single species ‘yield-perrecruit’ model of fish stocks. The output (yield) is controlled by adjusting the mortality and size of fish that are caught. Instruments for limiting fishing mortality are catch quotas (TAC, total allowable catch) and limits on fishing effort. Since it is much easier to define and measure catch than effort, the former is more widely used. In many shared, international fisheries, such as those governed by the European Union, the annual allocation of catch quotas is the main instrument of fisheries management. This requires costly annual assessment of many fish stocks, which must be added to the operating costs when looking at the economic balance for a fishery. Because annual assessments are costly, only

the most important species, which tend to be less vulnerable, are assessed. For example the US National Marine Fisheries Service estimate that the status of 64% of the stocks in their area of responsibility is unknown. The instruments for limiting the size of fish caught are mesh sizes, minimum landing sizes and various kinds of escape panels in the fishing gear. These instruments can be quite effective, particularly where catches are dominated by a single species. In multispecies fisheries, which catch species with different growth patterns, they are less effective because the optimal mesh size for one species is not optimal for all. There are two classes of economic instruments for fisheries management: (1) property rights and (2) corrective taxes and subsidies. The former demands less detailed information than the latter and may also lead to greater stakeholder participation in the management process, because it fosters a sense of ownership. The management of demersal fisheries will always be a complex problem, because the marine ecosystem is complicated and subject to change as the global environment changes. Management will always be based on incomplete information and understanding and imperfect management tools. A critical step towards better management would be to monitor performance in relation to the target objectives and provide feedback in order to improve the system. One of the changes which has taken place over the past few years is the adoption of the precautionary


96


approach. This seeks to evaluate the quality of the evidence, so that a cautious strategy is adopted when the evidence is weak. Whereas in the past such balance of evidence arguments were sometimes applied in order to avoid taking management action unless the evidence was strong (in order to avoid possible unnecessary disruption to the fishing industry), the presumption now is that in case of doubt it is the fish stocks rather than the short-term interests of the fishing industry which should be protected. This is a very significant change in attitude, which gives some grounds for optimism in the continuing struggle to achieve sustainable fisheries and healthy ecosystems.

Demersal Fisheries by Region The demersal fisheries of the world vary greatly in their history, fishing methods, principal species and management regimes and it is not possible to review all of these here. Instead one heavily exploited area with a long history (New England) and one less heavily exploited area with a short history (the south-west Atlantic) will be described.

continued. Most groundfish species remain at low levels and, even with management measures intended to rebuild the stocks, are likely to take more than a decade to recover. The changes in abundance of ‘traditional’ groundfish stocks (cod, haddock, redfish, winter flounder, yellowtail flounder) have to some extent been offset by increases in other species, including some elasmobranchs (sharks and rays). However, prolonged high levels of fishing have resulted in severe declines in the less resilient species of both elasmobranchs (e.g. barndoor skate) and teleosts (halibut, redfish). This is covered more fully elsewhere (see Ecosystem Effects of Fishing). The second half of the twentieth century saw major changes in the species composition of the US demersal fisheries. By the last decade of the century catches of the two top fish species during the decade 1950–59, redfish and haddock, had declined to 0.6% and 2.3% of their previous levels, respectively. Shellfish had become the main element of the demersal catch (Table 3).

New England Demersal Fisheries

Demersal Fisheries of the South-West Atlantic (FAO Statistical Area 41)

The groundfish resources of New England have been exploited for over 400 years and made an enormous contribution to the economic and cultural development of the USA since the time of the first European settlements. Until the early twentieth century, large fleets of schooners sailed from New England ports to fish, mainly for cod, from Cape Cod to the Grand Banks. The first steam-powered otter trawlers started to operate in 1906 and the introduction of better handling, preservation and distribution changed the market for fish and the species composition of the catch. Haddock became the principal target and their landings increased to over 100 000 t by the late 1920s. The advent of steam trawling raised concern about discarding and about the damage to the seabed and to benthic organisms. From the early 1960s fishing fleets from European and Asian countries began to take an increasing share of the groundfish resources off New England. The total groundfish landings rose from 200 000 t to 760 000 t between 1960 and 1965. This resulted in a steep decline in groundfish abundance and in 1970 a quota management scheme was introduced under the International Commission for Northwest Atlantic Fisheries (ICNAF). Extended jurisdiction ended the activities of distant water fleets, but was quickly followed by an expansion of the US fleet, so that although the period 1974–78 saw an increase in groundfish abundance, the decline subsequently

The catch from demersal fisheries in the SW Atlantic increased steadily from under 90 000 tonnes in 1950 to over 1.2 Mt in 1998 (Figure 7). The demersal catch consists mainly of fish species, of which Argentine hake has been predominant throughout the 50-year record. The catch of shrimps, prawns, lobsters and crabs is almost 100 000 t per year and they have a relatively high market value. Thirty countries have taken part in demersal fisheries in this area, the principal ones being Argentina, Brazil, and Uruguay, with a substantial and continuing component of East European effort. A three year ‘pulse’ of trawling by the USSR fleet resulted in a catch over 500 000 t of Argentine hake and over 100 000 t of demersal percomorphs in 1967. The fisheries in this area are mostly industrialized and long range. As the stocks on the continental shelf have become fully exploited, the fisheries have extended into deeper water, where they take pink cusk eel and Patagonian toothfish. The main coastal demersal species are whitemouth croaker, Argentine croaker and weakfishes. The Argentine hake fishery extends over most of the Patagonian shelf. It is now fully exploited and possibly even overexploited. Southern blue whiting and Patagonian grenadier are also close to full exploitation. Hake and other demersal species are regulated by annual TAC and minimum mesh size regulations.



Table 3

Average US catch from the north-west Atlantic region

Species

Scientific name

Atlantic redfishes nei Haddock Silver hake Atlantic cod Scup Saithe (Pollock) Yellowtail flounder Winter flounder Dogfish sharks nei Raja rays nei American angler American sea scallop Northern quahog (Hard clam) Atlantic surf clam Blue crab American lobster Sand gaper Ocean quahog

Average catch (t)

Sebastes spp Melanogrammus aeglefinus Merluccius bilinearis Gadus morhua Stenotomus chrysops Pollachius virens Limanda ferruginea Pseudopleuronectes americanus Squalidae Raja spp. Lophius americanus Placopecten magellanicus Mercenaria mercenaria Spisula solidissima Callinectes sapidus Homarus americanus Mya arenaria Arctica islandica

1950–59

1989–98

77 923 64 489 47 770 18 748 17 904 11 127 9103 7849 502 73 41 77 650 52 038 43 923 28 285 12 325 12 324 1139

518 1487 16 469 24 112 3846 6059 5085 5700 17 655 10 304 19 136 81 320 25 607 160 795 57 955 29 829 7739 179 312

then the catches from some stocks have declined, due to overfishing, while other previously underexploited stocks have increased their yields. The limits of biological production have probably been reached in many areas and careful management is needed in order to maintain the fisheries and to protect the ecosystems which support them.

1.4 1.2 1.0

Million tonnes

97

0.8 0.6 0.4

See also

0.2 0.0

50

19

60

19

70 19 Year

80

19

90

Ecosystem Effects of Fishing. Fishing Methods and Fishing Fleets.

19

Other demersal

Weakfishes nei

Patagonian grenadier

Whitemouth croaker

Southern blue whiting

Argentine hake

Figure 7 Total demersal fish landings from the south-west Atlantic (FAO Statistical area 41).

Conclusions Demersal fisheries have been a major source of protein for people all over the world for thousands of years. World catches increased rapidly during the first three-quarters of the twentieth century. Since

Further Reading Cochrane KL (2000) Reconciling sustainability, economic efficiency and equity in fisheries: the one that got away? Fish and Fisheries 1: 3--21. Cushing DH (1996) Towards a Science of Recruitment in Fish Populations. Ecology Institute, D-21385 Oldendorf/ Luhe, Germany. Gulland JA (1988) Fish Population Dynamics: The Implications for Management, 2nd edn Chichester: John Wiley. Kurlansky M (1997) Cod. A Biography of the Fish That Changed the World. London: Jonathan Cape.


DETERMINATION OF PAST SEA SURFACE TEMPERATURES M. Kucera, Eberhard Karls Universita¨t Tu¨bingen, Tu¨bingen, Germany & 2009 Elsevier Ltd. All rights reserved.

by the CLIMAP (Climate: Long Range Investigation Mapping and Prediction) group (Figure 1), the recognition of the pattern, rate, and magnitude of ocean cooling and warming during glacial cycles, and the determination of tropical and polar temperatures during supergreenhouse climates of the Mesozoic and early Cenozoic.

Introduction Temperature is one of the most striking aspects of the climate system. Its distribution on the surface of the Earth determines regional climate and habitability and its changes through time mirror the climatic evolution of our planet. The surface ocean represents the main reservoir of latent heat with ocean currents being the major means of heat redistribution at the Earth’s surface. In addition to the position of currents, sea surface temperature also reflects vertical mixing in the ocean, where upwelling brings deep cold waters to the surface. Clearly, knowledge of past variations in sea surface temperature (SST) and its spatial distribution is essential for the assessment of current warming trends as well as for the general understanding of the dynamic processes in the ocean and their links with global climate. Importantly, quantitative data on spatial and temporal distribution of past SST are instrumental for validation of numerical climate models. Unfortunately, direct instrumental measurements of SST are only available for the last two centuries. This time window is too short to assess the nature of natural variability and understand how the ocean and atmosphere systems behave under different climatic regimes. To know when, why, by how much, and how fast SST changes, scientists need long, continuous records spanning centuries to millions of years. Fortunately, SST affects a range of metabolic and thermodynamic processes that leave a distinct signature in biological materials. Some of these materials are preserved in the geological record and the signals locked in them can be decoded to estimate SST variation in the past. Unsurprisingly, reconstructing the temperature history of the surface ocean became one of the major tasks in geosciences and the last 50 years have seen the establishment of quantitative paleothermometry as the backbone of paleoceanography and paleoclimatology. Among the main achievements of SST paleothermometry are the reconstruction of the ocean surface during the Last Glacial Maximum

98

Characteristics of Sea Surface Temperature Before explaining the various paleothermometry methods, it is important to understand what is actually meant by SST. The surface ocean derives its temperature from solar insolation, whereas the temperature of the deep waters is relatively constant, preserving the SST signature at the place of their origin. At present, deep waters in the oceans are formed chiefly in polar and subpolar regions and their temperatures range between 0 and 4 1C. The resulting thermal gradient (Figure 2) implies that only a relatively shallow upper layer of the ocean retains the temperature of the ocean’s surface. The ‘mixed layer’ is homogenized by wave action and turbulence. Its depth is typically 50–100 m (Figure 3), although in summer during calm weather and due to intense heating, it can be only a few meters thin. Importantly, the mixed layer can be taken to roughly correspond to the photic zone, which means that phytoplankton and photosymbiotic organisms like corals and some planktonic foraminifera are likely to record the SST, whereas some zooplankton may live below the mixed layer, recording the temperature of the thermocline rather than the surface ocean (Figure 2). Due to the high thermal capacity of water, the range of surface ocean temperature is smaller than that of the atmosphere. At present, the coldest surface waters in polar oceans reach the freezing point of seawater at –1.7 1C while the highest surface temperatures of 32 1C are recorded in the Western Pacific Warm Pool. Seasonal variations in SST are also relatively muted, which makes the SST a good recorder of the mean state of the climate system. However, the magnitude of the seasonal signal (Figure 3) exceeds the precision of most methods for past SST determination and the difficulty in seasonal attribution of SST reconstructions is one of the main caveats in paleothermometry.


DETERMINATION OF PAST SEA SURFACE TEMPERATURES

99

Sea ice 75° N Land ice

Sea ice (seasonal)

50° 25° 0 25° 50°

Sea ice 75° S

Land ice 150° W

100°

50°

−7.5

50°

0

−5.0

−2.5

0.0

100°

150° E

2.5

Temperature change (°C) Figure 1 The CLIMAP reconstruction of SST during the Last Glacial Maximum (c. 21 000 years ago), expressed as the difference between present-day and glacial values. The map represents one of the milestones of paleoceanography. It was derived from abundances of microfossil species in sediment cores using the transfer function method. Reproduced from Mix A, Bard E, and Schneider R (2001) Environmental processes of the ice age: Land, oceans, glaciers (EPILOG). Quaternary Science Reviews 20: 627–657, with permission from Elsevier. Data from CLIMAP Project Members (1976) The surface of the ice-age Earth. Science 191: 1131–1137.

Recorders of SST signature

Idealized SST profile

Coral reefs

∼∼ ∼∼

25 °C

100 m 150 m ∼∼ ∼∼ Zooplankton

Geological archives of past sea surface temperature

15

Mixed layer

50 m Phytoplankton

Photic zone

Hermatypic corals

5

Thermocline

∼∼ ∼∼ ∼∼ ∼∼

2000 m 3000 m 4000 m

Deep-sea sediments Figure 2 An idealized scheme of the habitats of SST signal carriers and the occurrence of geological archives of SST. Hermatypic corals and phytoplankton grow only in the photic zone, which normally overlaps with the mixed layer. Some zooplankton species live below the photic zone and record thermocline temperatures rather than SST.


100


Average yearly SST variation off Puerto Rico 29.0 0m 28.5

10 m 20 m 30 m

28.0

50 m SST (°C)

75 m 27.5

27.0

26.5

26.0

25.5 Jan.

Feb.

Mar.

Apr.

May

Jun.

Jul.

Aug.

Sep.

Oct.

Nov.

Dec.

Month of year Figure 3 A typical tropical mean annual SST record for different depths. The temperature of the Caribbean Sea off Puerto Rico varies by 42 1C throughout the year; the thickness of the mixed layer varies between 450 m in winter and c. 30 m in summer. Data from the Comprehensive Ocean–Atmosphere Data Set (COADS).

Methods for Determination of Past Sea Surface Temperatures Nature of the Sea Surface Temperature Signal in Geological Materials

Many properties of marine organisms are affected by their chemical and physical environment. For SST reconstructions, the most significant are kinetic processes that alter the rate of incorporation of trace elements and isotopes into biominerals, metabolic regulation of the composition of cellular membranes, and temperature-driven changes in species abundance (Table 1). These various signals locked in fossil materials are not directly related to individual environmental parameters. Therefore, a whole branch of paleoceanography has been established to develop and test proxies – recipes and algorithms describing ways of how to relate measurements and observations made on fossils and other geological material to past environmental variables. Paleothermometry is the branch of paleoceanography devoted to the development of proxies for the reconstruction of past ocean temperatures. At present, only two archives have the capacity to provide long and continuous records of past SST: marine sediments and coral colonies (Figure 2). Deep-

sea sediments are particularly suitable as SST archives because they are often undisturbed and accumulate continuously for millions of years. Marine sediments are sampled by coring or drilling and SST signals are extracted from mineral or organic remains of planktonic microorganisms that lived in the surface ocean above the site of deposition. Strictly speaking, the SST signal is not in situ, because the plankton remains were transported to the seafloor by sinking through several kilometers of overlaying water. The degree of potential lateral displacement of the signal depends on the size of the signal carrier: small organic-walled microfossils and organic compounds are more easily transported than shells of planktonic foraminifera (Figure 4) or microfossils carried down in fecal pellets. SST records in marine sediments have typically centennial and rarely decadal resolution. The skeletons of hermatypic corals grow by accretion and like tree rings they record with high resolution the conditions of the surface ocean. Hermatypic corals are restricted to the tropical oceans with normal salinity where water temperature exceeds 25 1C. The vast majority of reef-building corals (such as the genus Porites) harbor photosymbiotic algae and are thus limited to the photic zone (Figure 2). Individual coral records recovered from cores drilled into living or fossil submerged coral heads are typically only a few



Table 1

101

Parameters of the main methods used to determine past sea surface temperatures

Method

Substrate

SST rangea

Typical standard error ( 1C)

Temporal applicability

Regional applicability

Oxygen isotopes in planktonic foraminifera Mg/Ca in planktonic foraminifera

Foraminiferal calcite

2 1C to no upper limit B5 1C to no upper limit

0.5b

0–120 Myc

All oceans

1

0–120 Myd

Mostly tropical waters, no upper limit Best performance between 5 and 27 1C 0–28 1C 428 1Ch

0.5b

0–130 ky

All oceans, best performance in Tropics Mainly Tropics

1.5

0–5 Mye

2

0–120 Myf

5–30 1C

1–1.5

0–500 kyg

Radiolaria

0–30 1C

1–1.5

0–500 kyg

Diatoms

2 to 18 1C

1–1.5

0–500 kyg

Dinoflagellate cysts

2 to 22 1C

1.5–2

0–500 kyg

Foraminiferal calcite

Oxygen isotopes and Sr/ Ca in corals

Sclearctinian aragonite

Alkenone unsaturation k0 ) (U37

Haptophyte alkenones in sediments Crenarchaeotal membrane lipids in sediments Planktonic foraminifera

Tetraether index (TEX86)

Transfer functions

All oceans except those of polar regions All oceans, possibly also lakes All oceans except those of polar regions All oceans except those of polar regions Polar to transitional waters All oceans except those of deep oligotrophic basins

a

Mean annual SST. Provided the isotopic and trace element composition of seawater is known. c In sediments older than c. 25 My in pristine preservation only. d In sediments older than c. 5 My, unknown species offsets and Mg content of the ocean may cause significant bias. e Ecology of pre-Quaternary alkenone-producing haptophytes is unknown. f Possibly older, if well-preserved lipids are found. g Estimate based on the age of the modern planktonic foraminifer fauna; older applications are problematic. h Several alternative calibrations are used for SST 4 28 1C b

centuries long. However, depending on the growth rate, they can provide yearly or even monthly resolution and data from several heads can be stacked to develop longer, regional records.

Paleothermometers Based on Chemical Signals in Biominerals

Oxygen isotopes in planktonic foraminifera Measurement of oxygen isotopic composition of marine biogenic carbonates provided geoscientists and paleoceanographers with the first quantitative paleothermometer. The method is based on the discovery by the American physicist and Nobel Prize laureate Harold Urey that during precipitation of calcite from seawater, a thermodynamic fractionation occurs between the two main stable isotopes of oxygen, 16O and 18O, and that this fractionation is a logarithmic function of temperature. The technique was originally calibrated on mollusk shells, but since

the pioneering work of Cesare Emiliani in the 1950s, planktonic foraminifera (Figure 4) became the prime target for oxygen isotope paleothermometry in oceanic sediments. The oxygen isotopic composition of foraminiferal shells is determined by mass spectrometry of CO2 evolved by acid digestion of foraminiferal calcite. Although modern analytical methods allow measurements of oxygen isotopes in single shells, a standard analysis requires about 10–20 specimens, in order to obtain a representative average of the conditions integrated in a fossil sample. The oxygen isotopic composition of foraminiferal calcite is expressed as a deviation, in per mille, from a standard (normally V-PDB): 2 6 d18 O ¼ 4

18

3 O=16 O 18 O=16 O standard 7 sample 5 103 18 O=16 O


standard

½1

102


Figure 4 Shells of planktonic foraminifera from a Caribbean surface sediment sample. Foraminifera are the main substrate for geochemical and transfer function paleothermometry. The shells have been extracted by washing and sieving of the sediments through a 0.15-mm mesh. The image is 4 mm across.

The determination of past SST from oxygen isotopes requires a priori knowledge of the isotopic composition of ambient seawater. When this is known, empirical calibrations from culture experiments or surface sediments can be used to derive paleotemperature equations for different species and SST ranges. These equations usually take on a secondorder polynomial form; the Shackleton calibration is perhaps the most commonly used: Tcalcification ½1C ¼ 16:9 4:38 d18 Ocalcite d18 Oseawater 2 þ 0:1 d18 Ocalcite d18 Oseawater ½2 Given the high analytical precision of oxygen isotope measurements (o0.1%), the method allows SST reconstruction with the precision of 0.5–1 1C. However, the biomineralization of calcite by foraminifera is affected by a range of secondary factors and the d18O composition of seawater in the past must be estimated from a number of assumptions.

The two main factors influencing d18Oseawater are the global ice volume and the local precipitation/ evaporation (P/E) balance. Due to fractionation during water phase transitions, continental ice is depleted in the heavy isotope, leaving the seawater isotopically heavy. Similarly, during evaporation, the light isotope preferentially enters the gaseous phase, making water vapor and precipitation isotopically light. The P/E effect is also known as the salinity effect, because seawater salinity is affected by the P/E balance in a similar way as its oxygen isotopic composition. The combined effect of ice volume and P/E balance on d18O in foraminifera can be substantially larger than that of SST, making paleotemperature reconstructions from foraminiferal d18O rather difficult. In fact, the d18O method is now used predominantly to reconstruct the isotopic composition of seawater and, if available, other paleothermometers are preferred. Planktonic foraminifera are heterotrophic zooplankton and their habitat is not limited to the surface mixed layer (Figure 2). Each species calcifies at specific depths, often below the mixed layer, and



Mg/Ca in planktonic foraminifera The amount of trace elements incorporated into inorganically precipitated calcite depends chiefly on their concentration in the solution from which the mineral grows. However, the rate of cation substitution in calcite also depends on temperature. This latter effect is particularly pronounced for magnesium and forms the basis of the Mg/Ca paleothermometer. The substitution of magnesium for calcium in calcite is endothermic and both thermodynamic calculations and experimental data confirm that Mg/Ca of inorganically precipitated calcite increases by 3% per 1C. A similar phenomenon is observed in biologically precipitated calcite produced by a range of organisms, although the actual relationship between Mg/Ca and SST is often offset from the thermodynamic prediction. In oceanic sediments, shells of planktonic foraminifera offer the best source of calcite for Mg/Ca paleothermometry. Planktonic foraminiferal shells are composed of low-magnesium calcite and contain several times less magnesium (o1% Mg) compared to inorganic calcite precipitated at the same conditions. The response of foraminiferal Mg/Ca to increasing SST is 9% per 1C, that is, 3 times larger than the thermodynamic prediction. The reason for the latter effect is not known, although it is likely to be related to the lower Mg/Ca content of foraminiferal calcite. No matter what its cause, the larger response of foraminiferal Mg/Ca to SST change is extremely important for paleothermometry as it increases the sensitivity of the method and reduces the uncertainty of Mg/Ca-derived SST reconstructions.

Following the thermodynamic prediction, the relationship between Mg/Ca in foraminiferal calcite and calcification temperature is expressed as an exponential function: h i Mg=Ca mmol mol1 ¼ b emTcalcification ½1C

½3

After a rigorous cleaning to remove Mg from adhering sediment, organic matter, and secondary precipitates, the magnesium content of foraminiferal shells is routinely determined by inductively coupled plasma mass spectrometry (ICP-MS) or atomic emission spectrometry (ICP-AES). Both techniques provide a precision of Mg/Ca ratio measurements of the order 0.5%. Data from laboratory cultures as well as calibrations from sediments and watercolumn samples converge on the value of m being between 10.7% and 8.8%. The pre-exponential term b appears more variable, ranging between 0.30 and 0.53. The parameters of the equation vary depending on which species is analyzed and how the calcification temperature is determined. An extensive calibration based on foraminifera from sediment traps (Figure 5) yielded the following conversion 6

5 Mg/Ca (mmol/mol)

maximum production occurs at specific times of the year. In addition, the presence of symbiotic algae in some species offsets the d18O signal, as do changes in growth rate, addition of gametogenic calcite at depth, carbonate ion concentration in seawater, and various other ‘vital effects’. Dissolution appears to increase d18O in foraminifera by about 0.2% per kilometer water depth, but this effect is relatively easy to contain. On geological scales, foraminiferal calcite is gradually recrystallized during burial in the sediment and the secondary calcite assumes the isotopic composition of inorganic precipitate from pore water fluid. Foraminiferal d18O paleothermometry in Cretaceous and Paleogene sediments appears only feasible on pristinely preserved shells recovered from clay-rich sediments. Although the method is no longer seen as the prime SST paleothermometer, the development of independent methods based on the same signal carrier (Mg/Ca, transfer functions) provides an exciting opportunity to subtract the SST contribution to foraminiferal d18O and reconstruct more precisely than before the d18O of seawater.

103

4

3

G. truncatulinoides G. trilobus G. sacculifer G. ruber (white) G. ruber (pink) P. obliquiloculata G. inflata G. hirsuta N. dutertrei G. crassaformis G. conglobatus

2

1 10

15

20 Temperature (°C)

25

30

Figure 5 The relationship between calcification temperature (inferred from oxygen isotopes) and Mg/Ca in a range of planktonic foraminifera species from a Bermuda sediment trap. For most species, the global calibration expressed in eqn [4] can be used, yielding a standard error of 1.5 1C. Modified from Barker S, Cacho I, Benway HM, and Tachikawa K (2005) Planktonic foraminiferal Mg/Ca as a proxy for past oceanic temperatures: A methodological overview and data compilation for the Last Glacial Maximum. Quaternary Science Reviews 24: 821–834. Data from Anand P, Elderfield H, and Conte MH (2003) Calibration of Mg/Ca thermometry in planktonic foraminifera from a sediment trap time series. Paleoceanography 18: 1050, with permission from Elsevier.


104


equation, which appears to be applicable globally and across a range of species: Tcalcification ½1C ¼

1 0:09

h i1 Mg=Ca mmol mol1 A ½4 ln@ 0:38 0

The standard error of foraminiferal Mg/Ca calibrations is typically around 1 1C. Because of the exponential nature of the conversion equations, the sensitivity of the Mg/Ca paleothermometer decreases with temperature. The method is thus less suitable for SST reconstructions in subpolar and polar waters, whereas its sensitivity is higher in the Tropics where it ought to yield robust reconstructions even when past tropical SST was higher than today. As discussed in the previous section, planktonic foraminifera migrate through the water column during life and different parts of the shell calcify at different depths, often within or even below the thermocline. In order to circumvent this problem, Mg/Ca paleothermometry generally focuses on surface-dwelling species, which spend most of their life within the mixed layer: Globigerinoides ruber, Globigerinoides sacculifer, and Globigerina bulloides. The sedimentary Mg/Ca signal is typically derived by a pooled analysis of 10–50 shells and is therefore biased toward the season of highest production of the analyzed species. The effect of salinity on foraminiferal Mg/Ca (for the range of normal marine conditions) is believed to correspond to 0.6–0.8 1C per psu; the effect of surface ocean pH is c. –0.6 1C per 0.1 pH units. Both these effects are relatively small and affect Mg/Ca SST reconstructions only in extreme environments or on geological timescales. On these timescales, the possibility of changes in oceanic Mg/Ca must be considered as well. By far the largest secondary effect on Mg in foraminiferal shells is that of carbonate dissolution. Mg is preferentially removed from calcite during dissolution, and in foraminifera sinking through the water column and on the seafloor, this effect translates to about 0.4–0.6 1C loss per km or 2.8 1C loss per km of lysocline shift. With some effort, the influence of carbonate dissolution on Mg/Ca SST reconstructions can be quantified and it is generally assumed that in normal marine sediments, the maximum bias on the three commonly used species is of the order of 0.5 1C. Foraminiferal Mg/Ca paleothermometry is analytically relatively straightforward and offers the possibility to determine past SST from the same

phase as oxygen isotopes. This provides a powerful means for subtracting the temperature signal obtained from Mg/Ca measurements from the d18O signal in foraminifera, allowing reconstructions of past seawater isotopic composition and assessment of phase relationships between SST and ice sheet dynamics in the past. The greatest advantage of the Mg/Ca method is its high sensitivity at warm SST with no upper limit to reconstructed SST values. Therefore, Mg/Ca in planktonic foraminifera is currently considered the best paleothermometer for the Tropics. Sr/Ca and oxygen isotopes in hermatypic corals Skeletons of hermatypic corals also offer a possibility to obtain long continuous records of past SST variability. Unlike planktonic foraminifera, scleractinian corals produce aragonite, which is the rhombohedric variety of calcium carbonate. Elemental and isotopic substitutions in aragonite are also controlled by temperature and both the d18O and Sr/Ca of coral skeletons record SST. The d18O signal in scleractinian aragonite follows the slope of the relationship derived from inorganic precipitation experiments, but the curve is offset from equilibrium. The offset is presumably due to biological ‘vital effects’ such as growth rate or symbiont activity, and it appears to vary even within individual coral colonies. In order to convert coral d18O to past SST, the d18O of seawater must be known. The combined influence of the vital effect and uncertainties in the determination of seawater d18O make the coral d18O signal very difficult to deconvolve and the signal is thus often interpreted as a generalized ‘climate’ record, rather than SST. The Sr content of inorganically precipitated aragonite is an exponential function of temperature, but the exponential constant is small (–0.45% per 1C), and within the range of marine tropical SST, the relationship is often described as a linear function. The substitution reaction is exothermic, favoring Sr substitution at lower SST. Coralline Sr/Ca is determined by atomic absorption spectrophotometry (AAS) and the standard error of empirical calibrations is around 0.5 1C. However, recent studies indicate that the Sr content of coralline aragonite also responds to a number of secondary effects including symbiont activity, and the slope of empirical calibrations shows a large variability. In addition, the Sr/Ca paleothermometer may be affected on glacial/interglacial and longer timescales by secular changes of the Sr content of seawater. In comparison with calcite, aragonite is much more susceptible to dissolution and alteration, especially when exposed to meteoric waters. Both the



d18O and Sr/Ca signals are heavily altered during this process and the applicability of coral paleothermometry on uplifted reefs is severely reduced. The best coral climatic records are therefore derived from submerged corals. Coral paleothermometry is obviously limited to the Tropics, where hermatypic corals occur, and it provides an important means of studying tropical climate variability. Due to the high resolution of the records, climate data extracted from corals have provided fascinating insights into decadal dynamics of tropical oceans. It is, however, not always possible to separate the SST component in the d18O and Sr/Ca signals in coral skeletons and the method is therefore more useful when seen as recording the combined effect of SST, P/E balance, and seawater chemistry. Paleothermometers Based on Organic Biomarkers in Sediments 0

k Alkenone unsaturation (U37 ) The alkenone k0 (Uk standing for ‘unsatuunsaturation index U37 rated ketones’) is the main organic biomarker proxy for SST. It is based on the measurement of the relative abundance in marine sediments of unbranched long-chained methyl ketones with 37 carbon atoms (C37) and a variable number of double bonds. C37 alkenones, together with C38 and C39 alkenones, their fatty acid methyl esters (alkenoates), and alkenes, are specific for certain species of haptophyte algae, including one of the main primary producers in the oceans, the bloomforming Emiliania huxleyi (Figure 6). These molecules are chemically inert and survive transport throughout the water column as well as burial in marine sediments, where they are known to persist for millions of years, in concentrations ranging

105

typically between 0.1 and 10 ppm. The abundance of the individual compounds is determined from total lipid extracts or purified lipid fractions by high-temperature elution gas chromatography. Each analysis requires between 1 and 10 g of dry sediment. Current technology allows determination of the unsaturation index in all marine sediments, except where virtually no organic matter is preserved. The function of the long-chained alkenones is not fully understood. The high abundance of these compounds in the haptophyte algae (up to 10% of cellular carbon) indicates that they play a significant role in the cellular structure. It has been suggested that alkenones may serve as storage molecules or that they may be incorporated into the cellular membrane where they act as fluidity regulators. The latter hypothesis links alkenone unsaturation directly with growth temperature and provides a mechanistic explanation for the striking relationship between the degree of unsaturation and temperature, observed in C37 alkenone extracts from surface sediment samples and from laboratory cultures. The degree of unsaturation has been originally expressed as a ratio between the quantities of the di- (C37:2), tri- (C37:3), and tetra- (C37:4) unsaturated C37 alkenones: k U37 ¼

½C37:2 C37:4 ½C37:2 þ C37:3 þ C37:4

½5

It was later shown that there is no benefit in including the C37:4 alkenone and the currently used version of the index is marked by a prime: 0

k ¼ U37

½C37:2 ½C37:2 þ C37:3

½6

O

C37:2 heptatriaconta-15E,22E-dien-2-one O C37:3 heptatriaconta-8E,15E,22E-trien-2-one O 2 µm

C37:4 heptatriaconta-8E,15E,22E,29E-tetraen-2-one

Figure 6 Scanning electron micrograph (SEM; in false color) of a single cell of the alkenone-producing haptophyte alga Emiliania huxleyi. The cell is covered by a mineral skeleton (coccosphere), consisting of isolated interlocking plates (coccoliths). The structures k0 paleotemperature index are shown to the of the three unsaturated alkenones produced by this species and used to calculate the U37 right. SEM image courtesy of Dr. Marcus Geisen and Dr. Claudia Sprengel, Alfred Wegener Institute for Polar and Marine Research, Bremerhaven, Germany.


106


1.0 0.9 0.8 0.7

K′ U37

0.6 Atlantic

0.5

Indian

0.4 Pacific

0.3 0.2

Atlantic Ocean Indian Ocean Pacific Ocean

0.1 0.0 0

5

10 15 20 25 Annual mean SST, 0 m (°C)

30

Figure 7 Empirical calibration of mean annual SST at 0-m k0 determined in 149 depth and the alkenone unsaturation index U37 globally distributed surface sediment samples. Thick line shows the global regression as expressed in eqn [7]; standard error of the regression is 1.5 1C. Reproduced from Mu¨ller PJ, Kirst G, Ruhland G, von Storch I, and Rosell-Mele´ A (1998) Calibration of k0 based on core-tops the alkenone paleotemperature index U37 from the eastern South Atlantic and the global ocean (601 N– 601 S). Geochimica et Cosmochimica Acta 62: 1757–1772, with permission from Elsevier.

This modified index shows a strong and consistent linear relationship with SST (Figure 7). Empirical calibrations in surface sediment samples yielded the following conversion equation between the unsaturation index and annual mean SST at 0-m depth: 0 k 0:044 =0:033 SST0 m ½1C ¼ U37

½7

The form of this equation has been confirmed by laboratory culture studies as well as measurements of particulate organic matter in the water column. The standard error of the empirical calibration is about 1.5 1C and the method appears to work best in the SST range between 5 and 27 1C. The alkenone content of marine sediments is the result of many years of integration of alkenone production over the site of deposition. Therefore, the signal should be biased toward the preferred depth and season of growth of the haptophyte algae. It has been repeatedly demonstrated that the alkenone unsaturation index measured in surface sediments shows a stronger correlation with annual mean SST at 0-m depth than for any other depth and season. Consequently, past SST reconstructions based on alkenone unsaturation in fossil sediments are

interpreted as reflecting annual mean values at the ocean surface. Like all other algae, haptophytes are dependent on photosynthesis and their growth occurs within the photic zone (Figure 2). Therefore, the alkenone unsaturation signal records the conditions within or near the mixed layer and this conjecture ought to remain valid for as long as the haptophytes were the sole producers of the alkenones. The main production of haptophyte algae occurs in spring or in autumn. The temperature of these seasons is comparable to the annual average, explaining the best fit of sedimentary unsaturation values with mean annual SST. However, in the absence of information on the ecology of the alkenone producers in the past, interpretations of alkenone unsaturation values in fossil sediments as reflecting past SST for a specific season remain speculative. The modern alkenoneproducing species evolved during the late Quaternary (within the last 1 My) and ecological shifts could be significant for the interpretation of older alkenone signals, particularly at high latitudes where seasonal differences in SST are large. Apart from changes to vertical and seasonal production of the alkenones, the unsaturation signal appears robust to the usual sources of bias. Alkenones are recalcitrant to diagenesis and reworking of older alkenones has been found significant only in rare circumstances, such as in the polar waters where the primary production by the haptophytes was low, particularly during glacial times. Similarly, there is little evidence that the signal is affected by other environmental parameters such as productivity or salinity (within the range of normal oceanic values). To conclude, the fortuitous discovery by the Bristol organic geochemistry group in 1986 of the strong relationship between alkenone unsaturation and SST has survived extensive scrutiny and the unsaturation index remains one of the most robust means for determination of past SST in marine sediments. Crenarchaeotal membrane lipids (TEX86) The relatively novel biomarker paleothermometer TEX86 is based on the abundance of different types of crenarchaeotal membrane lipids called glycerol dialkyl glycerol tetraethers (GDGTs) in sediments. The presence of cyclopentane rings in the GDGT lipids of hyperthermophilic Crenarchaeota increases the thermal transition point of the membrane (the temperature beyond which the membrane loses its biological functionality) and the number of such rings in crenarchaeotal GDGT is known from laboratory cultures to increase with temperature. The kingdom Crenarchaeota of the domain Archaea



was long believed to include exclusively hyperthermophilic organisms, but recent studies showed that one group of Crenarchaeota is abundant in the upper water column of the oceans, where it may make up a significant portion of the picoplankton. Oceanic crenarchaeota also synthesize GDGT lipids

107

and the proportions of four types of GDGT molecules with 86 carbon atoms and different numbers of cyclopentane rings correlates with SST (Figure 8). Empirical calibration of GDGT abundance in surface sediments led to the definition of the TEX86

HO O O O

O

OH

I HO O O

O

O OH

II HO O O

O

O III

OH

0.90 0.80

TEX86

0.70 0.60 0.50 0.40 0.30 0.20

0

5

10

15 20 Annual mean SST (°C)

25

30

Figure 8 Empirical calibration of mean annual SST at 0 m depth and the crenarchaeotal membrane lipid index TEX86 determined in 40 globally distributed surface sediment samples. The thick lines show the linear regression expressed in eqn [9]; the standard error is 2 1C. The structure of the GDGTs is shown; roman numerals refer to the number of cyclopentane rings. The structure of the GDGTIV is not yet known. Modified from Schouten S, Hopmans EC, SchefuX E, and Sinninghe Damste´ JS (2002) Distributional variations in marine crenarchaeotal membrane lipids: A new tool for reconstructing ancient sea water temperatures? Earth and Planetary Science Letters 204: 265–274, with permission from Elsevier.


108


‘tetraether index of lipids with 86 carbon atoms’: TEX86 ¼

ð½GDGTII þ ½GDGTIII þ ½GDGTIV0 Þ ð½GDGTI þ ½GDGTII þ ½GDGTIII þ ½GDGTIV0 Þ ½8

This index is highly correlated with annual mean SST at 0-m depth and for SST range 0–28 1C the following conversion can be used: SST0 m ½1C ¼ ðTEX86 0:28Þ=0:015

½9

The abundance of the different GDGT types (Figure 8) is determined from total lipid extracts by highperformance liquid chromatography/atmospheric pressure positive ion chemical ionization mass spectrometry (HPLC/APCI-MS). The combined uncertainty in TEX86 SST reconstructions due to the standard error of the calibration and the analytical error is less than 2 1C. The TEX86 technique was first described in 2002 and there are many aspects of the method that require further investigation. The effects of secondary parameters like salinity and productivity are not known and it is not clear why the index correlates best with annual mean SST at 0-m depth, when the GDGT producing Crenarchaeota are known to inhabit the top 500 m of the water column. The abundance of marine Crenarchaeota is negatively correlated with that of algal phytoplankton and the GDGT lipids deposited on the seafloor should represent different seasons than those when algal blooms occur. On the other hand, the application of the TEX86 technique is not restricted to the oceanic environments. First results indicate that a similar relationship between SST and GDGT abundance also holds for lake sediments. Like alkenones, the TEX86 proxy can be used in carbonate-free sediments, and it additionally offers the possibility of reconstructing SST in very old sediments. The GDGT-producing Crenarchaeota are most likely ancient and measurable amounts of the GDGT lipids have been found in sediments older than 100 My. Paleothermometers Based on the Composition of Fossil Assemblages

Principles of transfer functions Temperature is one of the most important factors controlling the distribution and abundance of species of marine phyto- and zooplankton (Figure 9). Assemblage composition in those groups of plankton that leave a fossil record can therefore be used as a means to reconstruct past SST variation. In order to obtain quantitative, calibrated

reconstructions of past SST, microfossil assemblages are extracted from surface sediments, abundance of individual species is determined, and the data are linked to a long-term average SST above the site of their deposition. The relationship between the assemblage composition and SST is described in the form of the so-called transfer functions. Transfer functions are empirically calibrated mathematical formulas or algorithms that serve to optimally extract the general relationship between microfossil assemblage composition in sediment samples and environmental conditions in the surface ocean. Assuming that SST changes in the past were manifested mainly through zonal redistribution of microfossil assemblages, a transfer function derived from modern sediments can then be used to convert assemblage composition data from fossil samples to SST reconstructions (Figure 10). Unfortunately, this assumption is not always met and many fossil samples yield microfossil assemblages without analogs in the modern ocean. Such assemblages may reflect secondary alteration or reworking as well as no-analog oceanographic conditions in the past. No-analog assemblages are typically easy to recognize, but their use for SST reconstructions is always problematic. The potential of the transfer function technique in providing calibrated paleoenvironmental reconstructions has been realized in 1976 by the CLIMAP project in the seminal paper on SSTs of the Last Glacial Maximum. Ever since, transfer functions have become a standard method in paleothermometry, applicable in virtually all marine sediments and extendable to variables other than SST. The calibration of a transfer function requires a large database of census counts from surface sediments, including typically 100–1000 samples with counts of 20–50 species. The data set may be global or regional. Like any other empirical calibration, transfer functions rely on a number of assumptions (Figure 10). Unlike geochemical paleothermometers, transfer functions are robust to mild diagenesis, require no cleaning or chemical extraction, and can be relatively cheap and fast. However, they rely on the exact knowledge of species ecology and their use is thus limited to the late Quaternary (B0.1–0.5 My). Theoretically, transfer functions can be calibrated to SST representing any depth and season. In most cases, a temperature characteristic of the mean annual conditions in the mixed layer is used, as well as the winter and summer seasonal averages. However, a simultaneous reconstruction of summer and winter SST by the transfer function method does not equate to the reconstruction of a seasonal contrast in the past. The simultaneous reconstruction of several SST representations is only possible because all SST



109

90 G. ruber 80

Relative abundance (%)

70

60

T. quinqueloba

50

40

30

20

10

0 0

4

8

12

16

20

24

28

Annual average sea surface temperature (°C)

Figure 9 Relative abundance of two species of planktonic foraminifera (Turborotalita quinqueloba and Globigerinoides ruber) vs. mean annual SST at 10-m depth in 863 surface sediment samples from the North Atlantic (101 S to 901 N). Thick lines show average abundance per 1 1C increment. Note the strong and distinctly different SST response of the two species. Scanning electron micrographs of the two species are not to scale. Data from the MARGO database (http://www.pangaea.de).

representations in the present-day ocean are extremely highly correlated to each other. Transfer functions can only produce reliable SST reconstructions within the limits of the calibration range. The method is therefore most sensitive in the middle of the SST range and weaker at the SST extremes. Prediction errors are estimated by validation where a portion of the calibration database is held back, the transfer function is then applied on this validation data set, and the SST ‘reconstructions’ are compared with the actual values. If the calibration data set is small, the leaving-one-out method may be used where one sample is held back at a time and a new transfer function is produced. The prediction error is then calculated as the average of the prediction errors for the individual samples. The form of the transfer function may vary considerably. A simple linear relationship is expressed by the weighted average method, where past SST is calculated as the weighted average of the abundances of n species: , n n X X ðpi SSTi Þ pi ½10 SST ¼ i¼1

i¼1

where pi is the proportion of species i and SSTi is the ‘optimal temperature’ for species i determined empirically from the calibration data set. The performance of the weighted average method has been improved by including partial least squares regression. The classical method by Imbrie and Kipp is mathematically related. It involves transformation of the census counts into a number of artificial ‘assemblages’ by means of factor analysis, followed by a multiple regression of the assemblage scores (and their cross products and squares) onto SST. The modern analog technique (MAT) is a variant of the k-nearest neighbor regression. This method searches the calibration data set for samples with assemblages that most resemble the fossil assemblage. To identify the best analogs in the calibration data set, the square chord distance measure is used: dij ¼

p h i2 X ðxik Þ1=2 ðxjk Þ1=2

½11

k¼1

where dij is the dissimilarity coefficient between sample i and sample j, xk is the percentage of species k in a sample, and p is the total number of species in


110


Empirically determined

SST

Long-term average of instrumental measurements

1 Empirical calibration

Census counts in surface sediments

Modern fauna /flora

Transfer function

2 3

Census counts in fossil samples

Fossil fauna /flora

Reconstructed

Mathematically derived

4

5

Past SST

Assumptions of paleotemperature transfer functions: 1 2 3 4 5

Species abundances in the calibration data set are systematically related to sea surface temperature (SST). SST is a significant determinant in the ecological system represented by the calibration data set or is at least strongly correlated to such a determinant. The mathematical method used to derive the transfer function provides an adequate model of the ecological system and yields a calibration with sufficient predictive power. Species in fossil samples all have modern counterparts and their ecological responses to SST have not changed significantly. The joint distribution of SST with other environmental parameters acting on the ecological system in the calibration data set is the same as in the fossil sample.

Figure 10 The principles of the transfer function paleothermometer, showing the five main assumptions on which the method is based.

the calibration data set. Past SST is then reconstructed from the physical properties recorded in the best modern analog samples:

SST ¼

m X

, ðsi SSTi Þ

i¼1

m X

si

½12

i¼1

where SSTi is the SST value in ith of the m best analog samples and si is a similarity coefficient between the fossil sample and the ith best analog sample:

sij ¼

p X pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi xik xjk

½13

k¼1

The MAT approach has been further modified in the SIMMAX technique which weighs the best analogs additionally by their geographical distance from the fossil sample. The revised analog method (RAM) expands the calibration data set by interpolation in the space of the reconstructed environmental parameters. It also includes an algorithm for flexible selection of the number of best analogs.

Increasingly more complex mathematical methods are being used to extract the relationship between species counts and SST, including artificial intelligence algorithms and Bayesian techniques. The main issue in further development of transfer functions is how to avoid overfitting and extract the general signal. A useful approach appears to be that of comparison of the differences among SST reconstructions derived by different mathematical methods. Microfossil groups used for transfer functions Four microfossil groups are commonly used for transfer function paleothermometry. By far the most commonly used are planktonic foraminifera (Figure 4). The foraminifera are isolated from marine sediments by washing and sieving and the determination of foraminiferal assemblage census typically involves counting under a binocular microscope of 300–500 specimens in random splits of the 40.150 mm fraction. Foraminiferal transfer functions are applicable in all marine sediments deposited above the calcite lysocline, apart from polar waters where the dominance of a single species (Neogloboquadrina pachyderma) causes lack of resolution (Figure 11).



111

Artificial neural network transfer function for North Atlantic planktonic foraminifera 30

Reconstructed mean annual SST10 m (°C)

25

20

15

10

5 R 2 = 0.99 RMSEP = 0.96 °C N = 863

0

−5 −5

0

5

10

15

20

25

30

Observed mean annual SST10 m (°C) Figure 11 Comparison of actual SST values with transfer function estimates for a calibration of census counts of 26 foraminiferal species in 863 North Atlantic (101 S to 901 N) surface sediment samples. The transfer function was derived by training of backpropagation artificial neural networks. The root mean square error of prediction of the calibration is 1 1C; the calibration breaks down in polar samples (SSTo3 1C) due to the dominance of a single species. Modified from Kucera M, Weinelt M, Kiefer T, et al. (2005) Reconstruction of sea-surface temperatures from assemblages of planktonic foraminifera: Multi-technique approach based on geographically constrained calibration datasets and its application to glacial Atlantic and Pacific Oceans. Quaternary Science Reviews 24: 951–998. Data from the MARGO database (http://www.pangaea.de).

Typically, annual, summer, and winter SST are reconstructed simultaneously. Both regional and global calibrations exist, all showing prediction error estimates in the range of 1–1.5 1C. The approach has been successfully used throughout the last 0.5 My; applications to sediments as old as the mid-Pliocene exist, but must be viewed with utmost caution due to evolutionary modifications of species ecology. In sediments affected by carbonate dissolution, in the tropical Pacific and in polar oceans, the siliceous skeletons of diatoms (Chrysophyceae) and radiolarians are the prime target for transfer functions. The siliceous skeletons of diatoms and radiolarians are concentrated using chemical maceration of the sediment. The residue is mounted on glass slides and typically 300–500 specimens are counted under light microscope. Radiolaria are eurybathyal heterotrophic protists and their production maximum occurs in tropical and subpolar/transitional waters. The

maximum production of the diatom algae is in polar waters and high-productivity upwelling regions. Diatoms are particularly suitable for SST reconstruction in polar waters of the Southern Hemisphere. As light-dependent phytoplankton, their production during polar winter ceases and diatom transfer functions are therefore commonly used to reconstruct summer SST only. Unlike foraminifera, diatoms are diverse even in the coldest waters and can be used to reconstruct SST up to the freezing point of seawater. In fact, diatom transfer functions can be even used to reconstruct the extent of permanent sea ice. The prediction error of diatom transfer functions are estimated to be around 1 1C; slightly higher values are reported for radiolarians. Both groups have been used for paleothermometry only in the late Quaternary (last 0.5 My). Cysts of dinoflagellate algae are made of resistant biopolymers and can be extracted from sediments by


112


chemical maceration involving hydrofluoric acid. The residue is mounted on glass slides and preferably 300, but sometimes less, specimens are counted. Dinoflagellates are normally phototrophic, although some species are mixotrophic or heterotrophic. Cystproducing forms live typically in coastal waters and the use of dinocyst transfer functions in open-ocean settings is therefore somewhat limited. Large calibration data sets exist only for the Northern Hemisphere and dinocyst transfer functions are particularly useful in the high latitudes. Like diatoms, dinoflagellates can be used to reconstruct SST up to the freezing point of seawater and even to detect the presence of permanent sea ice. The prediction error of dinocyst transfer functions are estimated at 1.5–2 1C and the method has been used only in the late Quaternary (last 0.5 My). Apart from the four microfossil groups mentioned above, coccolithophore algae have been used for transfer functions in the past, but due to taxonomic uncertainties, this method has been largely abandoned. The transfer function approach has seen a recent revival, fueled mainly by the development and application of advanced computational techniques. Despite its caveats and limitation, the method is less sensitive to shifts in production season of the microfossils, is taphonomically more robust, and offers a methodologically independent validation of geochemical paleothermometers.

Conclusion The methods used to reconstruct past SST depend on analytical instrumental precision and computational power and their development has accelerated in the past few decades hand in hand with the technological progress. After the first attempts to quantify the magnitude of past SST change, the science is now focusing on the reduction of uncertainties in SST reconstructions, on a better understanding of the origin of the SST signal and on the search for new types of SST signatures. Existing methods are being tested on new substrates, such as oxygen isotopes in diatom opal and trace elements in monospecific isolates of coccolithophore calcite. Modern instrumentation is opening up new chemical and isotopic systems and the analysis of calcium isotopes in planktonic foraminifera appears a promising avenue of research. Finally, physical properties of microfossils other than assemblage composition are being studied and it appears that size of planktonic foraminifera and morphology of coccoliths (such as Gephyrocapsa oceanica) could be used to reconstruct past SST.

See also Calcium Carbonates. Land–Sea Global Transfers. Past Climate from Corals. Plankton and Climate. Protozoa, Planktonic Foraminifera. Radiative Transfer in the Ocean. Satellite Remote Sensing of Sea Surface Temperatures. Sediment Chronologies. Sedimentary Record, Reconstruction of Productivity from the.

Further Reading Anand P, Elderfield H, and Conte MH (2003) Calibration of Mg/Ca thermometry in planktonic foraminifera from a sediment trap time series. Paleoceanography 18: 1050. Bard E (2001) Comparison of alkenone estimates with other paleotemperature proxies. Geochemistry Geophysics Geosystems 2 (doi:10.1029/2000GC000050). Barker S, Cacho I, Benway HM, and Tachikawa K (2005) Planktonic foraminiferal Mg/Ca as a proxy for past oceanic temperatures: A methodological overview and data compilation for the Last Glacial Maximum. Quaternary Science Reviews 24: 821--834. Bemis BE, Spero HJ, Bijma J, and Lea DW (1998) Reevaluation of the oxygen isotopic composition of planktonic foraminifera: Experimental results and revised paleotemperature equations. Paleoceanography 13: 150--160. Birks HJB (1995) Quantitative palaeoenvironmental reconstructions. In: Maddy D and Brew JS (eds.) Technical Guide 5: Statistical Modelling of Quaternary Science Data, pp. 161--254. Cambridge, MA: Quaternary Research Association. CLIMAP, Project Members (1976) The surface of the iceage Earth. Science 191: 1131--1137. de Vernal A, Eynaud F, Henry M, et al. (2005) Reconstruction of sea-surface conditions at middle to high latitudes of the Northern Hemisphere during the Last Glacial Maximum (LGM) based on dinoflagellate cyst assemblages. Quaternary Science Reviews 24: 897--924. Gagan MK, Ayliffe LK, Beck JW, et al. (2000) New views of tropical paleoclimates from corals. Quaternary Science Reviews 19: 45--64. Gersonde R, Crosta X, Abelmann A, and Armand L (2005) Sea surface temperature and sea ice distribution of the Southern Ocean at the EPILOG Last Glacial Maximum – a circum-Antarctic view based on siliceous microfossil records. Quaternary Science Reviews 24: 869--896. Herbert TD (2006) Alkenone paleotemperature determinations. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6: The Oceans and Marine Geochemistry, pp. 391--432. Oxford, UK: Elsevier. Kucera M, Weinelt M, Kiefer T, et al. (2005) Reconstruction of sea-surface temperatures from assemblages of planktonic foraminifera: Multi-technique approach based on geographically constrained calibration datasets and its



application to glacial Atlantic and Pacific Oceans. Quaternary Science Reviews 24: 951--998. Lea DW (2006) Elemental and isotopic proxies of past ocean temperatures. In: Elderfield H (ed.) Treatise on Geochemistry, Vol. 6: The Oceans and Marine Geochemistry, pp. 365--390. Oxford, UK: Elsevier. Mix A, Bard E, and Schneider R (2001) Environmental processes of the ice age: Land, oceans, glaciers (EPILOG). Quaternary Science Reviews 20: 627--657. Mu¨ller PJ, Kirst G, Ruhland G, von Storch I, and RosellMele´ A (1998) Calibration of the alkenone paleotempek0 rature index U37 based on core-tops from the eastern South Atlantic and the global ocean (601 N–601 S). Geochimica et Cosmochimica Acta 62: 1757--1772. Schouten S, Hopmans EC, SchefuX E, and Sinninghe Damste´ JS (2002) Distributional variations in marine crenarchaeotal membrane lipids: A new tool for

113

reconstructing ancient sea water temperatures? Earth and Planetary Science Letters 204: 265--274. Waelbroeck C, Mulitza S, Spero H, Dokken T, Kiefer T, and Cortijo E (2005) A global compilation of late Holocene planktonic foraminiferal d18O: Relationship between surface water temperature and d18O. Quaternary Science Reviews 24: 853--868.

Relevant Website http://margo.pangaea.de – MARGO at PANGAEA; the MARGO project houses a database of geochemical and microfossil proxy results for the Last Glacial Maximum, numerous calibration data sets, software tools, and publications.


DIFFERENTIAL DIFFUSION A. E. Gargett, Old Dominion University, Norfolk, VA, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction Because three-dimensional turbulence normally occurs on time and space scales much smaller than those resolved by numerical ocean models, its effects must generally be parametrized in these models. A particularly important quantity is r0 w0 , the averaged (overbar) vertical flux of density associated with small-scale turbulence, in which r0 and w0 are turbulent fluctuations of density and vertical velocity. By analogy with molecular diffusive fluxes, this turbulent flux is routinely parametrized as r0 w0 ¼ Kr r¯ z , that is, as proportional to r¯ z , the vertical gradient of mean density with a proportionality constant, the turbulent eddy diffusivity Kr, that is assumed to be much larger than the molecular diffusivity of density Dr (see Three-Dimensional (3D) Turbulence). In addition, present ocean general circulation models (OGCMs) normally assume that the same diffusivity can be used to parametrize the vertical turbulent flux of all other scalars in terms of their mean gradients. This assumption underlies use of the same eddy diffusivity in separate equations for temperature (T) and salinity (S), the two properties that (with pressure) determine the density of seawater. Finally, frequent usage of constant Kr in OGCMs implicitly assumes that Kr will not change if mean ocean conditions such as stratification change over time. Details of parametrizations of small-scale turbulence would be a matter of relatively minor import were OGCM predictions insensitive to them, but this is not the case. Different constant values of Kr are known to produce large changes in important features of predicted steady-state circulations: for example, heat flux carried by the oceanic meridional overturning circulation, a climatically important variable, is particularly sensitive to the value chosen for Kr. OGCM sensitivity to Kr results from linkages among the vertical turbulent flux of density, the mean vertical density structure, and horizontal circulation. Since density in the ocean is a highly nonlinear function of the two scalar variables T and S, differences in their vertical fluxes will lead to changes in these modeled linkages, hence to potential changes in modeled circulation.

114

Differential vertical transfers of T and S have long been recognized in double diffusive processes (see Double-Diffusive Convection), which can occur in regions of the ocean where the mean vertical gradient of either T or S is destabilizing. However, even where mean gradients of T and S are both stabilizing, ordinary turbulent processes may be associated with what has come to be called differential diffusion, the preferential vertical turbulent transfer of T relative to S.

What Is Differential Diffusion? Differential diffusion results from the larger molecular diffusivity of T relative to that of S, by a mechanism whose cartoon is shown in Figure 1. Figure 1(a) shows a box filled with a light layer of warm (W) and fresh (F) water atop a denser layer of cold (C) and salty (S) water: both T and S components contribute to static stability of the system. Figure 1(b) depicts a blob of lower layer fluid that has been displaced into the upper layer: the blob becomes warmer relatively quickly as heat is rapidly communicated by molecular diffusion from its warmer surroundings. However, because the molecular diffusivity of salt (DS) is 100 times smaller than that of temperature (DT), very little salt escapes from the blob during the time it takes for its temperature to equilibrate. This leaves (Figure 1(c)) a blob that is still denser than its surroundings due to the excess salt, hence falls back to an equilibrium position (Figure 1(d)) between the two original layers. In this final state, temperature has been transferred vertically, but very little salt has accompanied it. In the case of actual turbulence, a simplified conceptual picture is that of a locally overturning turbulent eddy that stirs an embedded scalar field. The decay timescale te of the eddy is influenced by the turbulent Reynolds number Re uc/n (where u and c are characteristic turbulent velocity and length scales

(a)

(b) WF WF

CS T

CS

(d)

(c) T T CS

WF WS

WF WS

CS

CS

Figure 1 Cartoon of the process of differential diffusion of T and S, which has its roots in the much larger molecular diffusivity of temperature than salt in seawater, where DT C 100 DS.


DIFFERENTIAL DIFFUSION

Laboratory Evidence for Differential Diffusion Differential diffusion was first demonstrated in the laboratory, where J.S. Turner used T and S in turn to produce a two-layer density stratification that was subsequently mixed away by grid-generated turbulence. Figure 2 shows (normalized) entrainment velocity ue, which quantifies the vertical transport of density, as a function of a Richardson number Ri0 g(Dr/r)c/u2 defined in terms of turbulent velocity (u) and length (c) scales and the density difference Dr between the two layers. At high values of Ri0, the entrainment velocity is consistently higher when Dr is produced by T than when the same density difference was produced by S. Effects similar to those observed in Turner’s singly stratified experiments were subsequently demonstrated in laboratory experiments where T and S made simultaneous and equal contributions to the stability of either a density step or linear density stratification. In all of these laboratory settings, the oceanically important scalars T and S, characterized by DTC100 DScDS, exhibit differential diffusion in the sense of enhanced vertical flux of T relative to

Limit at zero Ri0 1.0 5×10−1 2×10−1 10−1 ue /u

and n is fluid kinematic viscosity), which determines the range of scales present in the flow, hence the time required to transfer turbulent kinetic energy to the small spatial scales at which velocity variance is dissipated. Turbulent stirring also transfers variance of an embedded scalar to a scalar dissipation scale that varies as D1/2, where D is the scalar molecular diffusivity (see Three-Dimensional (3D) Turbulence). Important oceanic scalars such as T, S, dissolved oxygen, and chemical nutrients are all characterized by scalar Schmidt number Sc n/D41, hence by scalar variance dissipation scales that are smaller than velocity variance dissipation scales (see ThreeDimensional (3D) Turbulence). Straining scalar variance to a smaller scalar dissipation scale requires additional time, and this added time increases as Sc increases. As a result, the amount of scalar variance that is dissipated during te depends upon Sc. For ScB1, scalar variance can be entirely erased within the decay timescale of the velocity field, while effectively nondiffusive scalars (Scc1) will experience almost no transfer of variance to their much smaller diffusive scales within the same period of time. Since irreversible scalar mixing is rooted in the elimination of scalar variance by intermingling at the molecular level, these differences can result in differential vertical diffusion of scalars which have different molecular diffusivities.

115

5×10−2 2×10−2 10−2 5×10−3 2×10−3 1.0

2

5

10 Ri0

20 l Δ =g u2

50

100

200

Figure 2 Comparison between directly measured entrainment velocity (equivalent to vertical flux) resulting from grid-stirring on one side of a density interface produced respectively by T (filled circles) and S (open circles). Mean stratification increases with Richardson number Ri0 , here calculated with turbulent length c and velocity u scales characteristic of the grid turbulence. Reproduced from Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press, with permission from Cambridge University Press.

that of S, that is, greater vertical diffusion of the scalar with the larger molecular diffusivity. In the doubly stratified experiments, where vertical gradients of T and S simultaneously make equal contributions to stability, the differential fluxes can be directly interpreted as KT4KS, that is, a larger turbulent diffusivity for T than for S.

Numerical Simulation of Differential Diffusion Much of what is known about the mechanism of differential diffusion in double stably stratified systems has come from direct numerical simulation (DNS), in which dependence of the process on molecular properties of scalars is provided by explicit resolution of the necessary dissipative scales. Even with modern computers, however, computational limitations do not presently allow three-dimensional resolution of the very small spatial scales associated with dissipation of true salinity variance. Thus to date, simulations have


116


been run instead with a variable ‘S’ that has a diffusivity only 10 times that of salt, that is, DT ¼ 10 D‘S’, rather than the appropriate value of DT ¼ 100 DS. While this computational pseudo-salt ‘S’ variable will continue to be referred to as S, it should be kept in mind that numerical results using ‘S’ underestimate the magnitude of differential diffusion of true salt relative to temperature. One of the major contributions of DNS to the investigation of differential diffusion has been documentation of the essential role played by smallscale restratification (counter-gradient flux) associated with the scalar of smaller molecular diffusivity. Figure 3 shows results from a DNS in which turbulence generated impulsively at t ¼ 0 stirs an initially linear density gradient made up equally by T and S. In this visualization, the initial contortions of T and S

T

S

isosurfaces situated originally at mid-depth in the box are identical. However, as time goes on, the effect of differences in molecular diffusivity can be seen in the faster disappearance of fluctuations in T, through its more effective molecular diffusion. The slower molecular diffusion of S leaves fragments of excess S that are progressively unsupported as turbulent kinetic energy dies away (see particularly panels 6 and 7). These heavy S fragments fall back toward a stable position (the restratification process), producing a counter-gradient flux and partially reversing the downgradient flux of S that occurred in the initial turbulent overturning process. Thus the net vertical flux of S is smaller than that of T. The essential role of stable stratification in producing the small-scale counter-gradient fluxes that lead to differential diffusion is emphasized in the

T

S

1

5

2

6

3

7

4

8

Figure 3 Increasing numbers indicate time evolution of T and S isosurfaces that are initially horizontal and located at the vertical midpoint of a computational box. The computational fluid is linearly stably stratified with equal contributions to density from T and S. Turbulence imposed impulsively at t ¼ 0 stirs both T and S fields as it decays. Molecular diffusion of T is 10 times larger than that of S, so T fluctuations disappear more rapidly, leaving unsupported fragments of anomalous S (panels 6 and 7) which cause countergradient flux (restratification) of S during part of the decay process . Reprinted from Gargett AE, Merryfield WJ, and Holloway G (2002) Direct numerical simulation of differential scalar diffusion in three-dimensional turbulence. Journal of Physical Oceanography 33: 1758–1782.



(a)

t >> 0

t >> 0

t=0

t=0

ST (b)

Tz Sz

ST

Unstratified

Stratified

40

z

20 0

−20 −40 −40

−20

0 x

20

40 −40

−20

0 x

20

40

Figure 4 (a) Schematic of differential mixing by unstratified and stratified turbulence follows evolution of a fluid element whose T content is represented by the shaded circle and S content by the unshaded region within. In the unstratified case, Lagrangian displacement increases monotonically, and net vertical transport of S exceeds that of the more diffusive (leakier) component T, leading to KS4KT. In the stratified case, a restoring buoyancy force starts to act on displaced particles once some T has preferentially diffused out of the fluid element. Lagrangian displacement attains a maximum and then decreases as the fluid partially restratifies, leading to KSoKT. (b) Vertical plane projection of trajectories of 11 Lagrangian particles tracked in DNS of turbulence in stratified and unstratified systems, showing the behaviors suggested in (a). Reproduced from Merryfield WJ (2005) Dependence of differential mixing on N and Rr. Journal of Physical Oceanography 35: 991–1003, with permission of the American Meteorological Society.

cartoon of Figure 4(a). The impact of the faster molecular diffusivity of T on a particle containing anomalies of both T and S depends upon whether the ambient fluid is unstratified (left) or stably stratified (right). In both cases the faster diffusing T leaks out into the surroundings during Lagrangian displacement of the particle. In the unstratified case, Lagrangian displacement is monotonic, that is, on average the particle continues to move in its initial direction, taken here as upward. In the position shown, the particle has lost much of its T anomaly and retained more of its S anomaly, resulting in greater vertical flux of S than T, hence KS4KT

117

in this unstratified case. In the stratified (oceanic) case, however, leakage of T enhances a restoring buoyancy force that eventually causes the Lagrangian displacement of the increasingly salt-heavy particle to reverse direction. Averaged over a turbulent event, the result is greater vertical flux of T, hence KSoKT. Figure 4(b) illustrates the fundamental differences in average Lagrangian displacements proposed in the cartoon, using Lagrangian particles tracked in DNS of stratified and unstratified turbulence. At an initial time, all the particles shown were situated near the x-axis and had upward initial velocities, as assumed in the cartoon of Figure 4(a). Net upward displacements are smaller in the stratified case, as a result of the restratification tendency associated with differential diffusion of T and S. Similar results are found in a DNS of differential diffusion of T and S ¼ ‘S’ associated with Kelvin– Helmholtz shear instability, thought to be the most common source of episodic turbulent mixing in the stratified interior of the ocean. Figure 5 is a visualization of the two scalar distributions, in terms of their (initially equal) contributions to density, during roll-up of the instability. Effects of the different dissipation scales of the two scalars are immediately obvious in the much finer spatial structure seen in rS relative to rT. While the fact that the distribution of rS contains much smaller scales does not itself necessarily imply differential diffusion, these simulations do exhibit differential diffusion: values of diffusivity ratio d KS/KT lie between 0.5 and 0.85. The density ratio Rr aT¯ z =ðbS¯ z Þ quantifies the relative contributions of T and S stratification to density stratification in a doubly stable system: extremes of Rr ¼ 0 and Rr ¼ N are respectively cases where either T or S is a passive scalar, that is, has no effect on density. Although initial DNS of differential diffusion used Rr ¼ 1, subsequent computations have investigated cases where T and S make unequal contributions to the basic density stratification. The impact of Rr on d proves to be systematic (d is smaller when Rr ¼ 0 than when Rr ¼ N) but secondary in importance to that of the strength of the turbulence: in all cases, d approaches 1 as the strength of turbulence increases.

Oceanic Values of Diffusivity Ratio The magnitude of flux differences that might actually be realized in the ocean is thus dependent upon the strength of typical ocean turbulence relative to the strength of turbulence in the DNS which are the main source of quantitative predictions. Unfortunately it is difficult or impossible to make


118


0

(a) y 0.5

z 0.5

0

ρS

(b) z 0.5

0

0

1

x

2

ρT

Figure 5 DNS results showing partial densities (a) rS and (b) rT during roll-up of a Kelvin–Helmholtz instability on a transitional layer between two water masses that are equally stably stratified by T and S. Values colored range from 0.4D (red) to 0.4D (dark blue), where 2D is the density difference across the transitional layer. Values outside this range are transparent. Reprinted from Smyth WD, Nash JD, and Moum JN (2005) Differential diffusion in breaking Kelvin–Helmholtz billows. Journal of Physical Oceanography 35: 1004–1022.

observational determinations of parameters such as turbulence Reynolds and Froude numbers that are typically used to characterize turbulence strength in DNS computations. Various determinations of diffusivity ratio d from laboratory, computational, and observational studies have instead been compiled as a function of Reb ¼ e/nN2, (where e is the rate of turbulent kinetic energy dissipation per unit mass 1=2 is the buoyancy frequency and N ¼ gr1 ¯z o r determined by the mean density stratification), a form of turbulence Reynolds number that is accessible to both computational and observational determination. Characterization in terms of Reb is doubly useful because Reb is also a commonly used metric of the strength of ocean turbulence: turbulent flows with Rebo200 are considered weak in the sense that a crucial characteristic of turbulence (isotropy) fails even at dissipation scales. Figure 6 shows a range of presently available data: most come from laboratory (thick curve, pluses) and DNS studies with Rr ¼ 1 (gray and solid circles, gray boxes), Rr ¼ 0 (white boxes), and Rr ¼ N (black boxes), but a single observational study has estimated d from microscale measurements of T and S,

albeit with large uncertainty represented by the large light gray circle. Recalling that the computational results (using ‘‘S’’ instead of S) underestimate the magnitude of differential diffusion for true salt, it is clear from Figure 6 that differential diffusion may become significant below values of order RebB1000. Since turbulence in the stratified ocean interior is most frequently characterized by Rebo1000, differential diffusion should act routinely to enhance the vertical diffusion of oceanic T relative to S.

Other Observational Evidence for Differential Diffusion? While the single set of observational determinations of d shown in Figure 6 results does suggest do1, it cannot be regarded as conclusive because d ¼ 1 is within observational uncertainty associated with the extreme difficulty of determining the S dissipation spectrum. However, other fragmentary pieces of observational evidence also point to the possibility that differential diffusion is active in the ocean interior.



119

1.1 1.0

−1.82

0.9

−1.84 T (°C)

d

0.8 0.7 0.6

−1.86 −1.88 −1.9

0.5 0.4

−1.92

0.3 0.2 10−1

100

101

102

103

104

34.6 34.605 34.61 34.615 34.62 34.625 34.63 S

Reb Figure 6 Diffusivity ratio d ¼ KS/KT as a function of buoyancy Reynolds number Reb ¼ e/nN 2. Estimates are from laboratory studies (thick curve: two-layer density stratification; pluses: linear density stratification), computational studies (gray bullets: impulsively generated turbulence; black bullets: turbulence generated by Kelvin–Helmholtz instability; squares ([black, gray, white] correspond to Rr ¼ [0,1,N]): impulsively generated turbulence), and ocean observations of T and S dissipation spectra (large light-gray circle).

One such piece of evidence may be found in the distinctive pattern of T/S fine structure found in overturning patches within doubly stable water columns. DNS results document the appearance of significant differences in turbulent scalar fluxes over timescales of order (0.1 0.2)TN, where TN ¼ 2p/N. Since the buoyancy period TN is the typical timescale for Kelvin–Helmholtz shear instability, the effects of differential diffusion should thus be evident within the lifetime of an overturning patch. Differential diffusion with do1 will result in preferential rotation of an initial T/S line toward horizontal, as weak turbulence preferentially mixes away fluctuations in the component with the larger molecular diffusivity. This sense of T/S rotation is seen in Figure 7 within an overturning event observed in a region of the Ross Sea, Antarctica, where mean T and S gradients are both stabilizing, enabling differential diffusion, and where T is effectively a passive scalar (i.e., RrB0), further enhancing the magnitude of differential diffusion of T relative to S. Preferential rotation in the sense seen in Figure 7 is a common feature of overturns observed in the Ross Sea. While both previous pieces of observational evidence for the action of differential diffusion in the ocean may be considered inconclusive, it is harder to dismiss observations of density-compensating T/S intrusions in a region that is doubly stable in T and S, that is, in a region where double diffusion cannot occur. Figure 8 shows a striking example, a set of

Figure 7 T/S plot through an overturn (red points) observed in a vertical profile from a region in the Ross Sea, Antarctica, where mean T and S gradients (blue line) are both stabilizing. Rotation of the T/S relationship toward horizontal, as observed in the T/S fine structure within the overturn, is a signature that would be expected if differential diffusion were significant in the sense expected from the molecular properties of T and S, i.e., with d ¼ KS /KTo1.

intrusive features in upper polar deep water, a water mass characterized by stable T and S stratification that is found between about 700 and 1000 m in the Arctic Ocean. Making necessary modifications to existing theory for intrusions driven by double diffusion, linear stability analysis determines values of KT and KS consistent with the observed intrusion thicknesses of 40–60 m. The required diffusivity ratio lies in the range 0.6pdp0.7, a not unreasonable value given the compilation seen in Figure 6. Moreover, required values of 1 10 6 m2 s 1p KTp 3 10 6 m2 s 1 are roughly an order of magnitude smaller than values determined from temperature measurements in the interior of other oceans, consistent with recognition that the Arctic interior is an environment of unusually low turbulence intensity; hence it is possibly a particularly favorable site in which to observe consequences of differential diffusion. Finally, with the predicted values of d and KT, growth rates for the intrusions are consistent with the appearance of measurable interleaving structures on timescales shorter than the estimated residence time of the upper polar deep water in the Arctic Ocean.

Does Differential Diffusion Matter? Although the molecular-scale mechanism of differential diffusion is similar to that of double diffusion, the energy supply for mixing is quite different in the two cases. While double-diffusive processes are


120


S (psu) 34.75 0

34.80

34.85

34.90

34.95

35.00

Diffusive convection Stable

p (dbar)

500

Fingering

Stable

1000

1500

Fingering 2000 −1.0

−0.5

0.5

0.0

1.0

1.5

T (°C) Figure 8 Within the Arctic Ocean, potential temperature generally increases with depth (pressure), while salinity exhibits a subsurface maximum, here seen B250 dbar. Horizontal lines bracket density-compensating intrusive features between about 700 and 1100 m, in doubly stable upper polar deep water. Differential diffusion with KT 4KS is a possible explanation for the existence of intrusions in a part of the water column that is stable to double diffusion. Reprinted from Merryfield WJ (2002) Intrusions in doublediffusively stable Arctic waters: Evidence for differential mixing? Journal of Physical Oceanography 32: 1452–1459.

associated with release of potential energy from the gravitationally unstable component of the mean T and S fields, differential diffusion is driven by the energy of ordinary small-scale three-dimensional turbulence. Thus while double-diffusive processes occur only where mean T and S gradients are suitable, differential diffusion will be active regardless of the signs of these gradients, provided only that the turbulent Reynolds number is sufficiently low. The compilation of d values shown in Figure 6 suggests that Rebo1000 may be sufficiently low. Since the majority of mixing events observed in the stratified interior of the ocean fall within this range, it appears that differential diffusion is potentially important for understanding net vertical density fluxes associated with three-dimensional turbulence in the ocean. Because changes in vertical density fluxes affect vertical density gradients (which in turn affect major ocean processes such as deep and bottom water formation, supply of nutrients to the bioactive surface layer, and ocean uptake of atmospheric carbon dioxide), and because such changes must be expected to evolve if density structure evolves, differential diffusion at microscales may have unexpected effects on much larger scales of ocean

variability. Investigation of the effects of parametrizations for double and differential diffusion found only small differences from results with equal constant diffusivities for T and S in a steady-state global ocean model. However, changes in relative vertical transports of T and S may allow previously unsuspected nonlinear feedback processes in ‘timedependent’ models, a possibility suggested by study of unequal diffusivities in a box model of the North Atlantic thermohaline circulation. With equal T and S diffusivities and fixed surface fluxes, the model exhibits the familiar result of convection in either the polar or subtropical set of boxes. However, with unequal diffusivities, there exists a range of atmospheric forcing under which the polar boxes instead oscillate between unstable convection and stable stratification, with a period that decreases with the value of the diffusivity ratio d ¼ KS/KTo1 in the polar boxes. While box models are certainly not the ocean, such results raise questions about unsuspected and unexplored feedback effects that may enter more realistic time-dependent numerical models, were they to properly incorporate the effects of differential (and double) diffusion. Given the importance of modelbased predictions of the evolution of the coupled atmosphere/ocean system under various climate forcing



scenarios, differential diffusion may yet prove to be more than a laboratory curiosity.

See also Double-Diffusive Convection. Estimates of Mixing. Three-Dimensional (3D) Turbulence.

Further Reading Gargett AE and Ferron B (1996) The effects of differential vertical diffusion of T and S in a box model of thermohaline circulation. Journal of Marine Research 54: 827--866. Gargett AE, Merryfield WJ, and Holloway G (2002) Direct numerical simulation of differential scalar diffusion in three-dimensional turbulence. Journal of Physical Oceanography 33: 1758--1782.

121

Jackson PR and Rehmann CR (2003) Laboratory measurements of differential diffusion in a diffusively stable, turbulent flow. Journal of Physical Oceanography 33: 1592--1603. Merryfield WJ (2002) Intrusions in double-diffusively stable Arctic waters: Evidence for differential mixing? Journal of Physical Oceanography 32: 1452--1459. Merryfield WJ (2005) Dependence of differential mixing on N and Rr. Journal of Physical Oceanography 35: 991--1003. Merryfield WJ, Holloway G, and Gargett AE (1999) A global ocean model with double-diffusive mixing. Journal of Physical Oceanography 29: 1124--1142. Nash JD and Moum JN (2002) Microstructure estimates of turbulent salinity flux and the dissipation spectrum of salinity. Journal of Physical Oceanography 32: 2312--2333. Smyth WD, Nash JD, and Moum JN (2005) Differential diffusion in breaking Kelvin–Helmholtz billows. Journal of Physical Oceanography 35: 1004--1022. Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press.


DISPERSION AND DIFFUSION IN THE DEEP OCEAN R. W. Schmitt and J. R. Ledwell, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 726–733, & 2001, Elsevier Ltd.

Introduction One of the very consistent results of oceanic microstructure measurements over the past two decades is that the rate of turbulent mixing is quite weak below the well-stirred surface layer. With few exceptions, the interior ocean seemed remarkably nonturbulent, indeed, almost laminar. A major puzzle arose, as the rate of renewal of deep waters seemed well in excess of what could be absorbed by turbulent vertical mixing in most of the oceanic gyres. Without mixing to warm the deep water, the large-scale meridional overturning circulation would cease, as this mixing is, in a very fundamental way, its driving mechanism. New observations in the abyssal ocean have shown that there are, in fact, sites of greatly enhanced turbulence, which appear to be sufficiently strong and extensive to provide the necessary mixing. The new observations, the basic physics involved, and the apparent energy sources for this turbulence are reviewed below.

The Thermohaline Circulation A major feature of the ocean circulation is the sinking of cold, dense water at high latitudes, its spread to low latitudes, and eventual upwelling. During this upwelling the water must be warmed and made less dense; this is accomplished by interior vertical mixing. This is a key step in the process, indeed it can be considered its driving agent, if the meridional temperature gradient is accepted as a given. The interior mixing can be thought of as a ‘suction’ which pulls cold water into the various basins; without this suction, the deep circulation would stagnate. Numerous modeling studies have now shown that it is the mixing rate that sets the strength of the thermohaline circulation. One of the long-standing puzzles in oceanography has been the apparent lack of sufficient mixing to accommodate the estimated production of cold bottom waters. Given the strength of bottom-water sources

122

and the area and stratification of the ocean basins, the required eddy diffusivity of turbulent vertical diffusion can be readily estimated; this turns out to be about 1 10 4 m2 s 1. However, numerous measurements by sensitive temperature and velocity probes of the microstructure of the ocean yield a typical midgyre value of diffusivity that is at least an order of magnitude smaller. This body of work has been confirmed by tracer release studies in the upper thermocline of the North Atlantic. Thus, the apparent lack of mixing cannot be ascribed to statistical sampling issues or the models used for the interpretation of microstructure data. Rather, it is now apparent that the main problem was one of lack of observations, since the majority of turbulence measurements were confined to the upper few hundred meters. This shows the vital importance of exploratory observations in the vast and undersampled oceans.

Deep-sea Observations of Mixing The clear importance of mixing to the thermohaline circulation has inspired recent attempts to sample the turbulent mixing rate in the deep sea. The variables of interest for estimating mixing rates are the dissipation rates of temperature variance and turbulent kinetic energy. These require measurements of centimeter-scale gradients of temperature and velocity with very sensitive instruments. There were a number of technical difficulties to overcome; sensors capable of detecting the subtle signatures of oceanic turbulence in the weakly stratified abyss can have difficulties withstanding the immense pressure at the bottom of the sea. Also, untethered instruments must be used, in order to reach 5–6 km depth and to avoid the vibrations introduced by trailing cables. This leads to a requirement for significant redundancy in tracking and weight-release mechanisms, to minimize the risk of losing a sophisticated instrument. The ‘High Resolution Profiler’ was developed at the Woods Hole Oceanographic Institution for studies of deep-ocean mixing. Profiles of dissipation below 3000 m depth began to be acquired in the early 1990s. The first were from the area around a submerged seamount in the eastern North Pacific. These showed substantial elevation of turbulence levels in the deep ocean near the seamount but there was no detectable enhancement of


DISPERSION AND DIFFUSION IN THE DEEP OCEAN

dissipation beyond 10 km away from the base of the seamount. Similar patterns were found in the Atlantic in a few deep dissipation measurements made in association with the North Atlantic Tracer Release Experiment. Profiles showed a rather uniform value of the vertical eddy diffusivity with depth except very near the bottom. These dissipation profiles were also the first to show that the models used to interpret turbulence measurements yielded mixing rates in good agreement with tracer dispersion in the upper thermocline (see Double-Diffusive Convection). These hints of enhanced mixing near topography helped to motivate a study of mixing in an abyssal fracture zone. Fracture zones are valleys that cut across midocean ridges, thus providing a passage for flow of cold bottom water from one ocean basin to another. The Romanche Fracture Zone on the Atlantic equator was suspected of harboring strong mixing because of the rapid change in water temperatures along the valley. That is, cold water entering the valley was observed to warm significantly as it flowed through. Turbulence observations confirmed that this was a site of intense mixing. The flow over the rough topography was strongly turbulent, though the mixing rates dropped dramatically above the layer of fast moving water. Though these observations provided the first solid evidence for strong mixing in the deep ocean, they were obtained in a rather specialized site, so that generalizations of mixing rates for an entire basin were not feasible. Thus, an effort was mounted to examine mixing over a more typical region of deep ocean bathymetry. The region next examined was in the Brazil Basin of the western South Atlantic. The Mid-Atlantic Ridge (MAR) poses a barrier to eastward movement of the coldest, densest bottom waters in the basin. Confined between the continental rise to the west and the MAR to the east, the Antarctic Bottom Water (AABW) enters the Brazil Basin through passages in the south and exits through the Romanche fracture zone and across the equator to the North Atlantic. Just as in the Romanche, the bottom waters are seen to warm as they progress through the basin. An estimate of the rate of input of cold water, and knowledge of the surface area of isotherms and vertical temperature gradients within the basin allow the required mixing intensity to be calculated. A number of investigators have made such estimates for the Brazil Basin, since the southern source flows and northern outflows were reasonably well measured. Similar to the earlier global estimates of deep water-formation rates, the required mixing coefficient is B1–4 10 4 m2 s 1.

123

Measurements of the turbulent dissipation rate made using the High Resolution Profiler revealed that the mixing was weak and insignificant in the western part of the basin. However, to the east approaching the MAR, there was a dramatic increase in turbulence. The bottom topography also changed character from smooth to rough from west to east, as can be seen in Figure 1. This basic picture of the localization of mixing over rough topography had been suggested by previous work but never before demonstrated. The pattern of turbulent dissipation suggested that a tracer release experiment near the MAR would be of greater interest than elsewhere in the basin. Accordingly, 110 kg of the tracer sulfur hexafluoride (SF6) were released at 4000 m depth at a point about 500 m above the ridge tops of the fracture zones that trend westward from the ridge crest. When sampled 14 and 26 months after the release, a dramatic spread of the tracer was observed both laterally (Figure 2) and vertically (Figure 3). The lateral dispersion is characterized by two features: the bulk of the tracer which drifts and spreads to the southwest, and a secondary maximum which propagates eastward toward the MAR. These two tracer lobes are especially apparent after 26 months. The rate of lateral spread is rather modest compared to similar experiments in the upper ocean, since the mesoscale eddies are weak at these depths. The east–west trending fracture-zone valleys, extending perpendicularly from the main north–south axis of the Ridge, are prominent features of Figure 2. These valleys play an important role in establishing the bimodal structure of the tracer plume. That is, the valleys seem to serve as conducts for the tracer that is carried to the east, a feature easily seen in the sections of Figure 3. It appears that the tracer mixed downward from the injection plume has been carried eastward toward the ridge, and tracer-free water has been advected in below the core of the plume. Propagation both along and across density surfaces is apparent. These sections were obtained in one of the fracture zones, where the tops of the confining ridges are at a depth of 4500 m in the vicinity of the main tracer plume. The diapycnal mixing coefficient estimates from the tracer dispersion is B3 10 4 m2 s 1 at the level of the injection and increases greatly toward the bottom, a pattern also displayed by the turbulence measurements. The separation of the plume into two main clouds by 26 months after injection suggests two mixing regimes are active at this site. The upper, interior, cloud seems to be spreading vertically, advecting southwestward, and sinking deeper across density surfaces. The downward motion is expected for a mixing rate


124


0 _ 500 _ 1000

Water depth (m)

_ 1500 _ 2000 _ 2500 _ 3000 _ 3500 _ 4000 _ 4500 _ 5000 _ 5500 _ 6000

_ 38 _ 36

_ 34

_ 32

_ 30

_ 28

_ 26

_ 22

_ 24

_ 20

_ 18

_ 16

_ 14

_ 12

Longitude

0.0

0.1

0.2

0.3

0.4

0.5

0.6

0.7 _4

2

0.8

0.9

2.0

5.0

8.0

22.0

_1

Diffusivity (10 m s )

Figure 1 A composite section of the vertical eddy diffusivity in the Brazil Basin, as estimated from direct turbulence measurements in the deep ocean. The rough bottom topography approaching the mid-Atlantic Ridge has a clear influence on mixing rates well into the water column above, whereas little mixing is observed above the smooth bottom region in the west.

that increases toward the bottom, though most oceanic models assume that upward flow is the rule. This finding means that the tracer experiment has added unique new knowledge to our understanding of the deep ocean. The flow of the lower tracer plume to the east is also exciting, as it shows how the deepest water is warmed in the course of eastward flow in the canyon. That is, the downward trend of the isopycnals requires a flow to the east in response to the pressure gradient; in steady state, the density gradient must be maintained by mixing (a steady state is a reasonable assumption, as the same density structure appears in surveys over 3 years). Thus, the fracture zones radiating from the ridge axis act as conduits which draw the bottom water toward the ridge, warming it and effectively upwelling it to lighter density strata. This is a secondary flow driven by the enhancement of mixing within the fracture zone valleys and toward the ridge. Such secondary flows just affect the stratification and circulation in the vicinity of rough bottom topography. These insights into the patterns of deep-ocean mixing are complemented by new information on the mechanisms causing the mixing. The overall pattern of spatial variation on the basin scale showed an

enhancement of mixing over rough topography. There was also an enhancement in the vertical shear in horizontal velocity on scales of B20–200 m in patterns that are indicative of internal waves. The variations in mixing can be further defined by examining the dissipation profiles above a variety of topographic types in the rough area. A simple classification of bottom types into crest, slope or valley profiles, and averaging of the data in a ‘height above bottom’ coordinate system, shows distinct differences between the average mixing rates (Figure 4). The ‘slope’ profiles show the greatest vertical extent of strong mixing above the bottom. This is to be expected if the bottom is a source of low-frequency internal waves or serves to reflect and amplify ambient internal waves. Internal waves propagate horizontally as well as vertically, with the lower frequency waves having a more horizontal propagation direction. Thus, the waves at a ‘slope’ station are likely to have come laterally from up-slope topographic features. The overall pattern of mixing variation has important consequences for estimating the net rate of mixing in the region. For topography to be a source of internal waves, there must be flow over the rough bottom.



125

_ 20

_ 21

Latitude

_ 22 _ 23

_ 24

_ 25 _ 26

_ 22

_ 20

_ 18

_ 16

_ 14

_ 12

_ 16

_ 14

_ 12

Longitude

(A) _ 20 _ 21

Latitude

_ 22 _ 23 _ 24 _ 25 _ 26

_ 22

_ 20

_ 18 Longitude

(B)

2000

2500

3000

3500

4000

4500

5000

5500

6000

Depth (m) Figure 2 Tracer distribution at 14 months (A) and 26 months (B) after injection. The red bars labeled ‘INJ’ mark the release site of the tracer. The contours denote the column integral of SF6 (in nmol m 2) and colors denote bottom depth. The station positions are shown as white dots; those with stars are used for the sections in the fracture zone valley of Figure 3. Southerly latitude and westerly longitude are shown as negative numbers.

This could be a mean flow, such as in the Romanche Fracture Zone, or a time-varying flow, due to eddies or the tides. Since the deep eddy field was not found to be particularly strong in this region, we suspect the tides are the main source of energy for the enhanced internal waves. A simple comparison of the 3-day averaged vertically integrated dissipation rate (in mW m 2) with the tidal velocities estimated from a global model is very suggestive of a dynamical link with the tides (Figure 5).

The record reflects the conditions at the geographical position of the dissipation profiles during the survey, and the tidal speed shows the amplitude of the estimated semidiurnal (12.42 hour period) tidal velocity over the spring–neap cycle during the cruise. The net dissipation is well correlated with the tides and appears to lag the forcing by about a day or two, a reasonable time scale for the vertical propagation of internal waves into the water column above.


126

DISPERSION AND DIFFUSION IN THE DEEP OCEAN 3500

Pressure (dbar)

4000

4500

5000

_ 21

_ 20

_ 19

_ 18

_ 17

_ 16

_ 15

_ 14

_ 13

_ 16

_ 15

_ 14

_ 13

Longitude

(A) 3500

Pressure (dbar)

4000

4500

_ 21

5000

_ 20

_ 19

_ 18

(B)

_ 17 Longitude

0

2

4

6 8 10 SF6 concentration (fM)

12

14

16

Figure 3 Sections of tracer concentration at 14 months (A) and 26 months (B) after release in an east–west oriented fracture zone of the Mid-Atlantic Ridge (MAR). The blue bar marks the injection site (INJ) and the white dots represent sample bottles. The white contours represent the potential density anomaly (kg m 3) referenced to 4000 dbar.

If one accounts for the variation in mixing rate due to the observed differences in slope and time of sampling relative to the tides, it is possible to estimate the net amount of mixing over the surveyed area, and make reasonable extrapolations for the rest of the Brazil Basin. When this is done, the mixing induced by tides and topography appears to account for the warming of Antarctic Bottom Water that is observed in the Brazil Basin. This apparent balance is

an important advance in our understanding of the deep flows in the ocean and the maintenance of the thermohaline circulation. However, it is too early to say whether tidally induced internal waves arising over rough topography generate enough mixing in the global ocean to absorb all the deep water production. Only a tiny fraction of the deep ocean has been surveyed with microstructure instruments. Models of how the tides



127

_

log10 ε (W kg 1) Valley profile _ 10 _ 9 _ 8

Crest profile _ 10 _ 9 _ 8

Slope profile _ 10 _ 9 _ 8

3500

z (m)

4000

Target isopycnal

4500

5000

5500 _ 23

_ 22.5

_ 22

_ 21.5

_ 21

Latitude Figure 4 Average profiles of turbulent kinetic energy dissipation rate above different types of topography. Profiles in a ‘height above bottom’ coordinate system are shown along a section of meridional topography at 18.51W, the longitude of tracer injection at the depth marked with a horizontal line. The dissipation is greater at the tracer level for sites above crests and slopes than over valleys.

4

7

5 2

4 3

1

2

_

3

6

Barotropic tidal speed (cm s 1)

_

Depth integrated dissipation (mW m 2)

8

1 0 440.0

0 450.0 460.0 Year day since 01 Jan 1996

470.0

Figure 5 Comparison of the vertically integrated turbulent dissipation rate and tidal speeds at the time and location of the profiles for the 1997 Brazil Basin cruise. A three-day running boxcar average has been applied to both variables.

interact with topography are only now being developed. Detailed bathymetric data and improved tidal models will be necessary to develop an estimate of the global mixing rate due to the tides. Consideration of the energetics of mixing for the thermohaline circulation suggests that the tides may cause about half the required mixing globally, making available about 3 mW m 2 on average. This is the minimum value of the column integrated dissipation rates observed in the rough regions of the Brazil Basin during the spring–neap cycle, though much lower values are obtained in smooth-bottom regions. The other potential source of enhanced turbulence in the deep ocean may be the flow of the Antarctic Circumpolar current over rough topography. This current receives a great deal of energy from its constant driving ‘tailwind’ and is a flow that extends to the bottom. The main dissipation mechanism is likely to be internal lee wave generation as the currents interact with complex topographic features. These waves can deliver energy to small vertical scales well up into the water column. Preliminary


128


data suggest that internal waves are indeed enhanced in rough areas of the Antarctic Circumpolar current, with an energy distribution suggestive of more uniform mixing in the vertical. This would provide a more conventional upwelling pattern than the complex circulations found in and around the valleys of the MAR. Estimates of the energetics involved suggest that it may be as important globally as the tides. However, no deep ocean turbulence measurements have been made in these remote areas, so the rates of mixing remain speculative. We must await the results of future scientific expeditions to explore this mechanism. Indeed, much also remains to be done on the tidal mixing issue, since the currently available data amount to little more than a glimpse into a complex problem.

Summary Whereas a decade ago, the mechanisms for mixing the ocean seemed quite mysterious, and the ‘missing mixing’ seemed to undermine our theories of the thermohaline circulation, we now have some leading candidates for how this mixing occurs. Bottom topography is key to these mixing mechanisms. They are as follows: 1. Flows through passages: fracture zones and other deep-sea passages serve to connect topographic deeps. Dense bottom waters can spill through these valleys at high velocity and experience strong mixing as they cascade over sills and rough topography. The change in deep-water properties has long been noted but now we know that the turbulence levels are indeed enhanced in such areas. 2. Tidal flows over rough topography: though midocean tidal velocities are weak, and thus generate negligible turbulence on their own, they readily interact with bottom topography to produce internal waves. These waves propagate into the water column above and produce enhanced turbulence to 1000–2000 m above the bottom. This mechanism predicts that interior mixing will be concentrated above rough bathymetry in areas with the strongest tides. It is not ‘boundary mixing’ in the traditional sense, since rough topography tends to be in association with midocean ridges and more heavily sedimented continental margins tend to have smooth bathymetry. The bottom source of energy for the waves leads to a decay of turbulence rates with height which leads to a general cross isopycnal flow ‘downward’. A diapycnal flow ‘upward’ is found in canyons, which serves to provide the requisite mixing and

upwelling for the bottom waters. This mechanism may be the leading mixing process serving to convert bottom waters into lighter density classes above midocean ridges throughout the world ocean. Antarctic Bottom Water in particular may be warmed largely by this process. 3. Mean flows over rough topography: this is a candidate mechanism about which little has been documented, but seems likely to play a role in key regions. The leading area of importance is the Southern Ocean, where the deep reaching Antarctic Circumpolar Current flows over significant bottom topography. Preliminary indications are that the vertical distribution of internal wave energy is more uniform, suggesting less variability in turbulent mixing rate, and thus a more uniform upwelling profile. This mixing may be key for converting the deep and intermediate waters into thermocline waters. The North Atlantic Deep Water is the primary candidate for warming in the region of the Antarctic Circumpolar Current.

These mechanisms all involve bottom topography, and thus point to the importance of improved knowledge of bathymetry for progress in understanding the deep and intermediate circulation. The spatial variations in mixing rates must lead to greater complexity in deep flows than has been anticipated in the present generation of models. It should also be noted that these mechanisms are not ‘boundary mixing’ in the usual sense. That is, the mixing occurs well away from the thin benthic boundary layer and is more likely to be concentrated over a midocean ridge than near a lateral boundary. Indeed, no enhancement of mixing has been observed in western boundary currents where the bottom is smooth. It is also important to note that numerous modeling studies have shown that the magnitude of the vertical mixing is limiting to the amplitude of the overall thermohaline circulation itself. Indeed, the role of interior mixing in the thermohaline circulation can be compared to the role of the wind stress in the wind-driven circulation; that is, turbulence provides the essential interior balance of vertical upwelling with downward mixing of heat, just as the wind stress pattern at the surface imparts circulation to the ocean’s horizontal gyres. The high latitude sinking regions are then analogs of the western boundary currents that close the wind-driven flows. This view more clearly shows that it is the interior mixing acting on available density gradients, rather than the surface formation of dense water, that acts as the driving agent for the thermohaline circulation. Indeed, without mixing, the



deep circulation would become cold and stagnant and oceanic warmth would be confined to a thin surface boundary layer. This issue is of major concern, since the substantial circulation of warm water poleward is responsible for much of the heat flux carried by the ocean. There is evidence that the North Atlantic limb of the thermohaline circulation was cut off at various times in the past, and some suggest that global warming could shut it off in future, due to surface water freshening by an enhanced hydrologic cycle. Recent modeling work shows that there is a delicate balance between the fresh water forcing and the rate of interior mixing that determines the stability of the thermohaline circulation. A better understanding of oceanic mixing is thus essential for prediction of the future evolution of the Earth’s climate system.

129

See also Double-Diffusive Convection. Internal Tidal Mixing. Tracer Release Experiments. Upper Ocean Mixing Processes.

Further Reading Ledwell JL, et al. (2000) Evidence for enhanced mixing over rough topography in the abyssal ocean. Nature 403: 179--182. Munk W and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research 45: 1977--2010. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96.


DISPERSION FROM HYDROTHERMAL VENTS K. R. Helfrich, Woods Hole Oceanographic Institution, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 733–741, & 2001, Elsevier Ltd.

Introduction Among the most significant scientific events of the last century is the discovery of hydrothermal vent fields and their unusual ecological communities along the crests of the mid-ocean ridges. The venting consists of localized sources of very hot (B3501C) water that rises 100–300 m above the vent before it spreads laterally, similar to the plume from a smokestack. Venting also occurs as less intense and relatively cool diffuse flow (B101C above the ambient ocean temperature) spread out over a much broader area than the focused high-temperature vents. Diffuse flow rises only a few meters above the seafloor before it is mixed with the ambient sea water. While the diffuse flow carries about half of the total hydrothermal heat flux, its effect on the overlying water column is much less dramatic than the high-temperature vents. The hydrothermal venting from diffuse and localized high-temperature venting is essentially continuous over periods of years to decades. On longer timescales the individual vent sites will dissipate and new sites will emerge at other locations along the ridge crest. This nearly continuous venting is also punctuated by intense short duration venting events. These intense events are produced by magma eruptions on the seafloor or tectonic activity that rapidly exposes large quantities of sea water to hot rock or releases large quantities of very hot water from the crust. The result is the creation of huge ‘megaplumes’ that can rise 500–1000 m above the ridge crest to form mesoscale eddies with diameters of O(20 km) and thickness of O(500 m). Because of its large buoyancy and the dynamical control exerted by the Earth’s rotation, vent fluid is not simply advected away by background flow. The venting is capable of forcing circulation on a variety of temporal and spatial scales and this may have important consequences on how the vent fluid is ultimately dispersed. This article focuses on the flow produced by high-temperature venting and megaplumes, since they are most relevant for long-range

130

dispersal of vent fluids as a consequence of their large vertical penetration into the water column. The fate of the heat, chemicals, and biological material released by the vent is of interest for many reasons. To geophysicists, the hydrothermal heat flux represents a substantial fraction of the total heat flux (conductive plus convective) from mid-ocean ridges. For chemists, the vent fluid is laden with chemicals and minerals leeched from the subseafloor rock that over geologic time may contribute to the geochemical state of the oceans. The unique biological communities that accompany venting depend upon the chemical and thermal energy delivered by the venting. Since most of these unusual animals can survive only at vent sites, the dispersal of vent fluid is the primary mechanism of larvae dispersal and the colonization of remote new vent sites.

The Rising Plume The cascade of scales initiated by a high-temperature vent begins with the fast O(1 h) rise of the buoyant fluid from the vent to the spreading level O(100 m) above the source. Fluid emerging from an isolated hot vent rises as a turbulent plume, entraining and mixing with the ambient sea water as it rises (see Figure 1). Because the entrained ambient water is denser than the fluid in the turbulent plume, the plume buoyancy decreases continually with height above the source. If the ambient environment had uniform density the plume fluid would remain less dense than the environment and it would rise indefinitely. However, even in the deep ocean the ambient water is stratified. Eventually the plume density increases until it equals the background density. After a short overshoot of this neutral density level due to the momentum of the rising fluid, the plume spreads horizontally as an intrusive density current, or it may be swept downstream by ambient currents. Figure 2 shows a transect through a hydrothermal plume on the Juan de Fuca Ridge in the North Pacific. The figure is typical of many such observations made worldwide over the last two decades. In the figure temperature and light attenuation anomalies (defined relative to the background values along isolines of density) are contoured as functions of depth and horizontal distance along the centerline of the axial valley. High values of light attenuation anomaly are due to particulates introduced into the water column by venting. Indeed, in many cases light attenuation is a more useful indicator of


DISPERSION FROM HYDROTHERMAL VENTS

Outflow

ZS ZM

Entrainment

Figure 1 Sketch of a plume from a localized high-temperature hydrothermal vent. The plume rises to a maximum height ZM above the source. The rising stem of the plume continually entrains ambient fluid so that the density of the plume buoyancy decreases with height above the bottom. Eventually the plume density equals the ambient density and it spreads laterally at some height ZS above the source.

hydrothermal activity than temperature anomaly. As discussed below, the temperature anomaly may be very small, or even negative. The main features of the buoyant rise, entrainment, and spreading processes can be determined with a theoretical plume model that conserves momentum, mass, and buoyancy integrated on a horizontal slice across the plume. The key assumption is that the entrainment velocity, or the rate at which ambient fluid is drawn into the plume, is linearly proportional to the vertical velocity within the plume. (Details of the basic plume models and the justification of the assumptions are discussed in Morton et al. (1956), see Further Reading section.) Modeling, laboratory experiments, and observations show that the maximum rise height the plume above the source ZM is given by: ZM ¼ 3:8 F0 N 3 where

1=4

r0 rs F0 ¼ Qg r0

½1

½2

and N2 ¼

gdra r0 dz

½3

131

Here F0 is the buoyancy flux from the vent and N is the buoyancy frequency of the ambient water, which over the rise height of the plume is assumed to be constant. Q is the source volume flux and rs (r0 ) is the source (ambient) fluid at the level of the vent. The background density is ra ðzÞ; z is the height above the source, and g is the acceleration due to gravity. While the maximum plume rise is ZM , the radial outflow is centered at a slightly lower height ZS E 0:8ZM . The thickness of the spreading layer over the source is E 0:2ZM . Typical values of F0 ¼ 102 m4 s3 and N ¼ 103 s1 give ZM E210 m. The time taken for a parcel of fluid to ascend from the vent to the spreading level BN 1 . Doubling the vent buoyancy flux leads to only a very minor change in ZM . This weak quarter power dependency on F0 is significant because observation of ZM and N are often used to estimate the heat flux from a vent H ¼ rs cp QðTs T0 Þ ¼ rs cp F0 =ga, where cp is the specific heat, a is the coefficient of thermal expansion and Ts and T0 are the temperatures of the source and ambient fluids, respectively. From eqn [1], HpZ4M N 3 . Estimates of H are very sensitive to small errors in either ZM or N. The model and eqn [1] were derived under ideal conditions. Others effects will affect plume behavior. For example, in an ambient flow with velocity U, ZM pðF0 U1 N 2 Þ1=3 . Increasing U leads to decreasing rise heights. Mid-ocean ridge crests are locations of rough, variable topography and this may affect plume behavior. For example, the slow spreading Mid-Atlantic Ridge is characterized by axial valleys that are typically deeper than the plume rise spreading level, ZS , while the fast spreading Pacific ridges have axial valleys shallower than ZS . Deep-valley topography will constrain the plume outflow and direct it along the ridge axis, limiting off-ridge dispersal of vent fluids. Despite limitations the basic plume model provides useful insight into the dispersal of vent fluids. The entrainment of ambient water into the plume causes a substantial dilution of a parcel of vent fluid. The volume flux into the spreading level QM ¼ 1:3ðF03 N 5 Þ1=4 . For F0 ¼ 102 m4 s3 and N ¼ 103 s1 , QM ¼ Oð102 m3 s1 Þ. This gives a dilution of Oð104 Þ for a typical source flux Q ¼ Oð102 m3 s1 Þ. Entertainment occurs at all levels, but the largest velocities of background fluid into the plume occur in the lower quarter of the rise height. Larvae of bottom dwelling vent organisms can easily be swept into the plume and rapidly transported up to the spreading level. They then have a greater likelihood of dispersal over the distances typical of individual vent spacing (O(10 km)), Furthermore, these larvae are in water that is chemically distinct


132


SLU 2, SSS Temperature anomaly (˚C) 1800

_

Light attenuation anomaly (m 1 )

1800 1900

1900 0.015 0.020

0.005 0.010

27.690 27.695

2000

0.025 0.030 0.035

2100

H 0.040

27.700

H

27.685

0.005 0.010 0.015

27.685

Depth (m)

Depth (m)

1 km

2000

27.690 27.695 0.050 0.030

27.700

0.025

2100

0.025

0.010 L

0.035

L

2200

2200 (A)

(B)

X Y

4.3 7.0

3.8 6.0

3.5 5.0

3.2 4.0

Figure 2 Transect of temperature (A) and light attenuation (B) anomalies through a hydrothermal plume on the Juan de Fuca Ridge in the North Pacific. The transect was taken along the axis of the axial valley. The maxima of temperature and light attenuation are located directly over the vent. (Reproduced with permission from Baker and Massoth, 1987).

from the ambient environment and this may enhance survival during the dispersal process. The temperature and salinity anomalies at the spreading level (where the density anomaly is zero) are dependent on the ambient temperature and salinity gradients and can be counter-intuitive. In the deep Pacific salinity decreases with height above the bottom, as does temperature. These background gradients results in relatively warm and salty spreading plume water. An example of the temperature and salinity vertical profiles through the effluent layer of a plume on the Juan de Fuca Ridge is shown in Figure 3. The spreading plume is easily distinguished as a layer of nearly uniform temperature and salinity in Figure 3A. In Figure 3B and C the potential temperature y and salinity are plotted against potential density, s2 , and clearly show the relatively warm and salty effluent layer. In comparison, in the deep Atlantic where the salinity increases with height above the bottom the spreading plume is relatively cold and fresh. The temperature of the plume at the neutral level is colder than the ambient water despite the enormous temperature of the source fluid. Thus temperature alone may not always be an obvious indicator of hydrothermal activity. In either case, temperature anomalies at the spreading level are Oð101 1CÞ despite source temperature anomalies of B3501C The rise characteristics of event megaplumes are similar to the continuous venting, except that the source duration is limited and the buoyancy flux, F0 , is typically one to two orders of magnitude larger. For comparison, the heat flux from a typical high temperature vent is 1–100 MW, while megaplume

sources are estimated to be >1000 MW. If the source duration is small compared with the parcel rise time, N1, then the plume model must be replaced by a model for an isolated thermal. In this case ZM ¼ 2:7ðB0 N 2 Þ1=4 , where B0 ¼ V 0 gðr0 rs =r0 Þ is the buoyancy and V0 the volume of the pulse of hot fluid forming the release. Entrainment into and dilution of a thermal are comparable to the continuous release.

Mesoscale Flow and Vortices A high temperature vent continuously delivers plume fluid to the spreading level. Ambient currents can simply advect this fluid away from the vent location, but if the currents are weak, or oscillatory with small mean, then plume fluid accumulates over the vent and a radial outflow must develop. On a timescale f 1 this radial flow will be retarded by rotation. Here f ¼ 2OsinðfÞ is the Coriolis parameter, O is the rotation rate of Earth and f the latitude. At 451N f ¼ 104 s1 . The outward-flowing fluid parcels turn to the right (looking from above in the northern hemisphere) and an anticyclonic circulation will develop. With time a slowly growing lens of plume fluid will be formed. Below the spreading level, entrainment into the rising limb of the plume causes a radial inflow of ambient water. The Coriolis acceleration again results in fluid parcels turning to the right as they move inward and cyclonic circulation is established. The result is a baroclinic vortex pair: an anticyclonic lens of plume fluid at the spreading level and cyclonic circulation of ambient fluid around the



1700

36.80

Density 2 36.84 36.88

36.92

34.54

Salinity (PSU) 34.58 34.62

34.66

1.8

Temperature (˚C) 2.2 2.0

2.4

Depth (m)

TT-175 STA.6

2000

2400 (A)

Effluent layer

S

T

r2

2.0

(˚C)

1.9

1.8

1.7

1.6 (B)

Salinit y (PSU)

34.62

34.60

34.58

34.56 36.84 (C)

2

36.88

36.92

Figure 3 (A) Vertical profiles of temperature, salinity and density through a hydrothermal effluent layer on the Juan de Fuca Ridge. The relatively warm and salty effluent layer is clear in plots of potential temperature, y, versus potential density, s2 (B) and salinity versus density (C). (Reproduced with permission from Lupton et al., 1985.)

133

rising buoyant plume. This circulation is sketched in Figure 4. The dynamical balance is geostrophic wherein the radial pressure gradients are balanced by the Coriolis acceleration. Figure 5 shows results from a laboratory experiment that illustrates the effects of rotation on plume structure. In the photographs dense fluid, dyed for visualization, is released from a small source into a tank of water that has been stratified with salt to give a constant density gradient (constant N). The tank is on a table rotating about the vertical axis to simulate the Coriolis effect. These photographs are taken looking in from the side a short time after the source has been turned on. The experiments were done with dense fluid which falls, rather than light fluid that rises. This is inconsequential for the physics and the hydrothermal vent situation can be envisioned simply by turning the figures upside down. As the rotation increases, as measured by decreasing values of the ratio N=f , lateral spreading of the plume is retarded and the anticyclonic lens of plume fluid becomes thicker. Dynamical scaling arguments and experiments show that the aspect ratio of the resulting eddy, h=L E 0:75f =N. Here h is the central thickness of the anticyclonic eddy (dyed fluid in the figure) and L is the radius. These arguments also give the eddy azimuthal, or swirling, velocity vBðF0 f Þ1=4. For the typical values of F0 and f ¼ 104 s1 , vB0:03 ms1 . This is comparable to observed background flows over ridges and suggests that plume vortex flow can persist in the presence of a background flow. The anticyclonic plume eddy will continue to grow until it reaches a critical radius LEZM N=f at which it becomes unstable and breaks up. An example of plume break up is shown in Figure 6, which contains a sequence of photographs looking down on the experiment. The plume vortex was initially circular (not shown), but eventually the eddy elongates (Figure 6A). It then splits into two separate vortex pairs (Figure 6B) which propagate away from the source (Figure 6C). The process of formation and instability process then begin again. A steady source results in the unsteady production of vent vortices as depicted in Figure 4. The timescale for this production process is tB B102 Nf 2 . For typical midlatitude values of N and f and ZM , LE2 km and tB B2 months. Note that f decreases as the equator is approached, resulting in larger diameter eddies which take longer to grow if other factors remain constant. The small size and long production time contribute to difficulty in directly observing these eddies, although observations of the water column


134


Anticyclonic circulation

Cyclonic circulation

Figure 4 Sketch of the effect of the Earth’s rotation on a hydrothermal plume. Rotation causes an anticyclonic horizontal circulation in the spreading fluid and cyclonic circulation below. These flows are indicated by the arrows. Lateral spreading of plume fluid is retarded by rotation and eventually the plume may become unstable, producing isolated vortices of plume fluid which have a radius LEZM N=f , which is about 2 km at mid-latitudes. A continuous vent could result in the production of numerous eddies which propagate away from the vent site.

properties do show indications of eddy-like features with the expected scales. Futhermore, ambient flows can be expected to influence this idealized scenario. But even with ambient flows that would tend to carry plume fluid from a vent, the tell-tale anticyclonic circulation at the spreading level and cyclonic flow below is expected. Indeed, there is observational evidence for this vorticity signature in time mean measurements of flow in the vicinity of a vent. However, the most compelling evidence for this dynamical scenario comes from megaplume observations. Figure 7 shows temperature and light attenuation (a measure of particulate concentration indicative of hydrothermal source fluid) anomaly sections across a megaplume observed near the Juan de Fuca Ridge in the North Pacific. Note the much larger rise height and lateral scales of this plume compared with the example in Figure 2. The structure of the megaplume is indicative of anticyclonic circulation within the core and this has been confirmed by detailed analysis. The production of eddies from either continuous high temperature venting or episodic megaplume

events is important for the dispersal of the vent fluid. While dispersal by simple advection and stirring by prevailing flows may be the dominant dispersal mechanism, even occasional eddy formation is significant. Coherent anticyclonic vortices are known to have closed streamlines and can retain their anomalous properties over long distances and large time periods. These eddies provide a mechanism for the long-range dispersal of vent organisms which are entrained into the rising plumes and then trapped in the eddies. Within the eddies larvae are suspended in water with anomalous properties that may enhance survival.

Large-scale flow Small-scale localized convection over the ridge crest can result in a large-scale circulation extending O(1000 km) from the ridge. From the point of view of the large-scale mid-depth (2000–3000 m) circulation, venting at numerous locations along a ridge crest segment produces an average net upwelling localized over the ridge crest characterized by



135

(A)

(B)

Zs Zmax

(C) Figure 5 Side-view photographs showing the effects of rotation on convective plumes. In (A) the rotation is zero. The classic turbulent plume and spreading layer are evident. Panel (B) has weak rotation, N=f ¼ 5:02. The lateral spreading is inhibited and the falling plume is partially obscured by the cyclonic circulation which has developed around the plume. In (C) the rotation is stronger, N=f ¼ 1:42, and the anticyclonic lens of dyed fluid is thicker and has a smaller radius. See the text for a description of the experiment. (Reproduced with permission from Helfrich and Battisti, 1991.)

divergent isopycnals over the ridge crest. This can set up a mean circulation similar to the circulation from an individual vent plume, anticyclonic flow at the spreading level and cyclonic below, but now extending along the length of the ridge crest segment. Fluid entrained into the plumes and upwelled to the spreading level must be replaced. This requires a broad downwelling flow to close the mass balance. However, on these scales of 100–1000 km the variation of the Corriolis parameter due to the spherical shape of the earth, the beffect, causes the two circulation cells to extend to the west of the ridge (regardless of hemisphere) to form what has been termed a beta;-plume. The ideal b-plume described

here will be affected by the ridge crest topography and any background mid-depth flow. However, there is some observational evidence suggestive of this model of long-range dispersal of plume fluid. Observations near 151S in the eastern Pacific (Figure 8) show a plume of anomalously high values of 3He (a distinctive signature of hydrothermal origin water) centered in the water column just above the depth of the ridge crest. The plume extends over 2000 km west of the ridge. As predicted by the b-plume dynamics the westward extension of the plume is greatest closer to the equator. There are no similar observation in the Atlantic; this is perhaps explained by the deep axial topography.


136


Discussion Localized high-temperature hydrothermal venting along ridge crest is capable of forcing circulations on scales many orders of magnitude larger than the vent field size. This is a consequence of the combination of the large buoyancy flux of hydrothermal vents and the dynamical effects of the Earth’s rotation. Rotating flows are very sensitive to vertical motions such that small vertical flows are amplified into large horizontal circulations. The immense buoyancy flux of the high-temperature vents and megaplumes gives rise to rapid vertical ascent and just as importantly large entrainment of background fluid into the rising plumes. These combine to force a localized net upwelling many times larger than the mass flux of the

(A)

1200

0.08 0.12

1400

Depth (m)

1600 1800 2000

Megaplume l Temperature anomaly

Profile

0.16

0.20 0.24

27.58

27.60 27.62

0.24

27.64

0.20 0.16 0.12 0.08 0.04 0.04

27.66

0.04 H

0.04

2200

0.04 H

2400 0

2

4

(B)

6

8 10 12 14 16 18 20 22 Distance (km) Megaplume l Attenuation

Profile

1200 1400

0.43 0.45 0.47

27.58

0.49 0.51

Depth (m)

1600 1800

27.60

0.51

27.62 0.49 0.47 0.45 0.43

27.64

2000

0.41 27.66

0.49 0.47 0.45

2200

0.43

0.41

H 0.41

2400 0

(C) Figure 6 Photographs of a laboratory experiment showing the formation and break up of a plume vortex. The view is from above and time increases from panel (A) to (C). A single continuous source produces one plume vortex which eventually becomes unstable, (A), and forms two smaller baroclinic vortex pairs which move away from the source (B), after which the process of plume vortex formation begins again (C). (Reproduced with permission from Helfrich and Battisti, 1991.)

2

4

6

8 10 12 14 16 18 20 22 Distance (km)

Figure 7 Observations of the temperature and light attenuation anomalies of a megaplume found near the Juan de Fuca Ridge. The figure shows a slice in depth and horizontal distance through the center of a nearly circular (in plan view) plume. The eddy aspect ratio b=LBf =N as predicted by the scaling theory and laboratory experiments. The lower level high in light attenuation may be the result of a separate, less intense hydrothermal source. The horizontal dashed lines are density isolines (sy contours) and the saw-tooth lines indicate the trajectory of the measurement package. (Reproduced with permission from Baker et al., 1989.)



137

3

STN.

7

( He)% 4

5

6

3

1

2

0 5 1

15

10 20

25 Depth (km)

2

3

4

35

30 45

40

30 25

50

20

5

1000 km

130˚

East Pacific Rise 0

120˚

1000 km

110˚ West longitude

100˚

90˚

Figure 8 Transect along 151S in the Pacific showing a plume of 3He anomaly, dð3 HeÞ, extending over 2000 km west of the East Pacific Rise. (Reproduced with permission from Lupton, 1995.)

individual vent. The stacked nature of the resulting horizontal flow, anticyclonic circulation at one level and cyclonic below, is typically unstable and produces eddies which have scales comparable to the local Rossby radius of deformation, Ld ¼ ZM N=f , based on the plume rise height. In reality the ultimate dispersion of high-temperature vent fluid probably occurs through a combination of simple advection and stirring by background flow and the formation of long-lived coherent vortices and b-plumes. The reader might wonder whether these rotationally influenced convective processes are at work in the atmosphere where smokestacks and fires routinely cause localized plumes. There is one important difference between the atmosphere and the ocean in this regard. The scale at which rotation influences the flow and would produce eddies, the deformation radius Ld , is very much larger in the atmosphere than the ocean due to the greater static stability of the atmosphere (larger N). So these features are not likely to occur as a consequence of smokestacks and fires, which are simply too small to be affected by rotation. However, hurricanes are an example of the interaction of convection and rotation which produces intense vortices. Also, it would be possible for large volcanic eruptions which rise into the stratosphere to produce the atmospheric equivalent of oceanic megaplumes. Finally, oceanic deep convection produced by surface cooling and sinking induces some of the same circulation characteristics discussed here, but over typically much larger horizontal scales than isolated high-temperature vents.

Nomenclature ZM Zs F0 N g Q r0 rs ra z y F H cp Ts T0 a s2 B0 V0 f O h L v tB Ld

maximum plume rise height. plume spreading level height. source buoyancy flux. background buoyancy frequency. gravitational acceleration. source volume flux. density of the ambient fluid at vent level. source fluid density. ambient density. depth above the source. latitude, potential temperature. latitude. heat flux. specific heat at constant pressure. source temperature. ambient temperature at vent level. coefficient of thermal expansion. measure of density. initial buoyancy of a thermal. initial volume of a thermal. Coriolis parameter. rotation rate of the Earth. vertical thickness of the anticyclonic plume eddy. radius of the anticyclonic plume eddy. azimuthal velocity within the plume eddy. timescale for plume break up. Rossby radius of deformation.


138


See also Hydrothermal Vent Biota. Hydrothermal Vent Ecology. Hydrothermal Vent Fluids, Chemistry of. Meddies and Sub-Surface Eddies. Mesoscale Eddies.

Further Reading Baker ET and Massoth GJ (1987) Characteristics of hydrothermal plumes from two vent fields on the Juan de Fuca Ridge, northeast Pacific Ocean. Earth and Planetary Science Letters 85: 59--73. Baker ET, Lavelle JW, and Feely RA (1989) Episodic venting of hydrothermal fluids from Juan de Fuca Ridge. Journal of Geophysical Research 94(B7): 9237--9250. Helfrich KR and Battisti T (1991) Experiments on baroclinic vortex shedding from hydrothermal plumes. Journal of Geophysical Research 96: 12511--12518. Humphries SE, Zierenberg RA, Mullineaux LS, and Thomson RE (eds.) (1995) Seafloor Hydrothermal

Systems, Physical, Chemical, Biological and Geological Interactions, Geophysical Monograph 91. Washington, DC: American Geophysical Union. Lupton JE (1995) Hydrothermal plumes: near and far field. In: Humphries SE, Zierenberg RA, Mullineaux LS, Thomson RE Seafloor Hydrothermal Systems, Physical, Chemical, Biological and Geological Interactions, pp. 317–346. Geophysical Monograph 91. Washington, DC: American Geophysical Union. Lupton JE, Delaney JR, Johnson HP, and Tivey MK (1985) Entrainment and vertical transport of deep-ocean water by buoyant hydrothermal plumes. Nature 316: 621--623. Morton BR, Taylor GI, and Turner JS (1956) Turbulent gravitational convection from maintained and instantaneous sources. Proceedings of the Royal Society of London, Series A 234: 1--13. Parsons LM, Walker CL, and Dixon DR (1995) Hydrothermal Vents and Processes. London: Geological Society.


DIVERSITY OF MARINE SPECIES P. V. R. Snelgrove, Memorial University of Newfoundland, Newfoundland, Canada Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 748–757, & 2001, Elsevier Ltd.

Overview The oceans comprise the largest habitat on Earth, both in area and in volume. Some 70.8% of the Earth’s surface is covered by sea water, making marine benthos (bottom-living organisms) the most widespread collection of organisms on the planet. Pelagic organisms (organisms that live in the water column above the bottom) occupy the sea water that fills the ocean basins, which represent some 99.5% of occupied habitat on Earth. Within these benthic and pelagic environments, there is a wide variety of habitat types from tropical to polar latitudes, that

ranges from the narrow band of rocky intertidal to open ocean surface waters to vast plains of muddy sediments on the deep-sea floor. Patterns of species composition and diversity vary considerably among habitats (Figure 1), although our understanding of these patterns is limited. Given the range and size of the habitats that occur in the oceans, this article will focus primarily on habitat and species diversity, while acknowledging the importance of genetic diversity. Within the oceans, the major variables that delimit distributions of organisms include tidal exposure, temperature, salinity, oxygen, light availability, productivity, biotic interactions, and pressure. For benthic organisms, substrate composition (rock, gravel, sand, mud, etc.) also plays a key role. Not all of these variables play the same role in different environments but the sum total of their interactions determines biological pattern in the oceans. Within the pelagic realm, organisms such as fish and whales are strong swimmers and can make

Decreasing biomass

Very low

Low _ Medium

Neritic estuarine

Oceanic surface

Depth (km)

3

5

9

200 _ 4000 m (slope)

Low _ High Reefs Seagrasses Kelps Subtidal sand Subtidal mud

Very low 7

Oceanic deep water 0 ± 20 m (intertidal) 0 _ 200 m (shelf)

0

1

Medium

4000 _ 6000 m (abyssal plain)

Very high Slope sediments

6000 _ 10 000 m (trenches)

Medium _ High

Rocky intertidal Mangels Salt marshes Sandflats Mudflats 300

Abyssal sediments

Very low Trenches

900 Approximate distance from shore (km)

Figure 1 Diagram indicating different marine habitats. In relative terms, the horizontal axis is greatly compressed; it should also be noted that some continental margins are much narrower than the example shown. Bottom habitats are shown in the lower left boxes and water column habitats are shown in boxes near the top of the figure. Above each habitat name, a relative estimate of diversity within the different habitats is shown, but these comparisons are only generalities for which there are important exceptions. Coral reefs are almost always highly diverse.


139

140

DIVERSITY OF MARINE SPECIES

significant headway against currents (nekton), whereas planktonic organisms, such as single celled and chain-forming algae (phytoplankton) and gelatinous forms (jellyfish, comb jellies, salps) drift passively with the predominant currents. Distributions of organisms are regulated by water mass characteristics such as salinity and temperature, water depth, and productivity; productivity varies with geography and depth in the water column. Life in the benthic environment is very different. Many benthic organisms have either limited mobility or are completely sessile. Much of their dispersal process occurs at the larval stage, which is planktonic for many species. The nature of the ocean bottom is a critical factor in determining the species that reside upon it. Species that attach to rocky substrate rarely occur in sediments, and species that occur in sands are usually relatively rare in mud. Shallow areas of the nearshore may be densely vegetated whereas deep-sea bottoms completely lack plant structure or photosynthetic primary producers of any sort.

The Organisms The number of described plant species in the oceans is relatively modest. Vascular plants, such as those found in seagrass beds, mangrove forests (mangels) and salt marshes, are few in number (B45 described species of seagrasses, for example), and the numbers of these species that co-occur at a given site are few. Described species of phytoplankton (B3500–4500) are considerably higher than for marine vascular plants, but by comparison with terrestrial plants (B250 000) the numbers are also modest. There has been some suggestion that phytoplankton taxonomists have tended to lump species, and that the number of described species in some groups is significantly less than the actual number of species, but this possibility has not yet been resolved. Defining species is even more problematic in marine microbes, a group for which it is believed that we have not yet even begun to describe their species numbers. In the past, species were defined based on characteristics of populations grown in culture. Recent advances in molecular biology have indicated that there are vast numbers of species that are completely missed with this approach, and that a small volume of sea water or sediment may contain tremendous microbial diversity. Invertebrates comprise most of the described diversity in the oceans. Some field researchers studying invertebrate communities subdivide the fauna into size groupings, rather than along strict taxonomic lines. Benthic ecologists define megafauna as animals

that can be identified from bottom photographs; macrofauna are those that are retained on a 300 mm sieve; meiofauna are those that are retained on a 44 mm sieve; and microbes are organisms that pass through a 44 mm sieve. Similar size groupings are used by plankton biologists. Examples of megafauna are fish and crabs, whereas macrofauna include polychaete worms, amphipod crustaceans, and small clams. Meiofauna include nematodes, small crustaceans, and foraminifera. Microbes include bacteria and protistans. In the past, larger organisms such as vertebrates have generated the greatest interest in terms of marine biodiversity and conservation, but in recent years the concern for smaller invertebrate and single-celled organisms has increased with the recognition that they form the bulk of the global species pool.

Sampling Biodiversity Organisms are subdivided on size out of practical necessity, in that the sampling approach and sample size that are appropriate for one group are often inappropriate for another. For megafauna living in the bottom or in the water column, trawls and nets of various sorts are used. In some cases, photographic identification is also used, but if a high degree of taxonomic resolution is desired then this approach is only useful for larger organisms. Similarly, acoustic devices use sound waves to estimate abundances of megafauna over comparatively large swaths of the ocean (hundreds of square kilometers can be covered in a day), but the taxonomic resolution is again poor and the approach has only limited utility for benthic habitats. For macrofaunal and meiofaunal groups, net samplers or pumps are used in the water column, and various types of grabs or corers are used to sample marine sediments. For microbial groups, a single cubic centimeter of sediment may typically contain 106 bacteria and it is clear that different sampling approaches are needed for different groups of organisms; typically, small water samples are collected for pelagic microbes whereas subsamples of bottom cores are typically taken for benthic microbes. The disparity in appropriate techniques for different size groups of organisms has contributed greatly to the paucity of studies on more than one taxonomic size grouping at a given locale. Unfortunately, where conflicting conclusions have been drawn about patterns in different groups of organisms, it is rarely possible to know whether the patterns truly vary among groups or merely reflect differences in sampling efforts.



Coastal Environments The shallowest marine environments are those forming the transition between land and sea, and include intertidal habitats that are regularly exposed to air, full sunlight, rapid changes in salinity and temperature and, at higher latitudes, ice abrasion. The best studied of these environments are rocky intertidal habitats, but sand- and mudflats, salt marshes, and mangels (mangrove communities) are other intertidal environments that occur at the land– sea interface and are alternately flooded with sea water, and then exposed to high temperatures, desiccation, and potentially hypersaline conditions. Intertidal habitats, because they are physically harsh environments that require specialized adaptation, are relatively low in diversity, although the modest numbers of species that are present are often represented by many individuals. Mangroves comprise globally only B50 species, and within a given area only one or two mangrove species may be present, but mangels contribute to local diversity patterns by creating habitat for marine and terrestrial species. Mangroves are limited to tropical waters, and in temperature and boreal latitudes a similar niche is filled by salt marshes. Marshes and mangroves are also extremely productive habitats, and although some of the detritus resulting from breakdown of the plant materials is exported to, and diluted in, the adjacent shelf system, decomposition of large amounts of this detrital organic matter can occur in bottom sediments within the marsh or mangel. As a general rule, near-shore environments that are highly productive are often relatively low in diversity. This is true of primary producers, which are often dominated by only a few species as occurs in marshes and mangroves, but also in near-shore areas where phytoplankton thrive under high nutrient conditions. The benthic communities in these environments are often low in diversity because a few species are able to take advantage of the abundance of food and are tolerant of the hypoxic conditions that are often associated with high organic input. All of these harsh environmental variables require specific adaptations for survival, and the relatively few species that are able to live under these conditions often occur in very high densities. This diversity is, nonetheless, important. The relative simplicity of intertidal habitats in terms of the numbers of species present and their accessibility has made them particularly conducive to experimental manipulation, and our understanding of biodiversity regulation in intertidal environments is probably greater than in any other marine habitat. Indeed,

141

intertidal systems provide a useful model community that has generated many of the major paradigms for marine ecosystems. Limited data from sedimentary communities suggest that paradigms developed in the rocky intertidal are not necessarily transferable to sedimentary habitats, but they do provide a useful framework of ideas for other environments. Intertidal habitats are important not only for the ecosystem services they provide (summarized below) but also because they represent a key ecotone (a transition zone between different communities) between land and ocean. The plants in mangels and salt marshes, for example, support a terrestrial fauna in their upper branches but a marine fauna in the sediments at their roots. These faunas may also interact with one another, and with other groups of organisms such as migratory birds that pass through the environment and feed while in transit. Estuaries represent another group of specialized coastal habitats, where fresh water and sea water meet and mix to form a gradient in salinity. Estuaries can encompass the intertidal habitats described above but they also contain a subtidal habitat and their influence may be felt well beyond the land–sea interface. Typically, estuaries are relatively low in diversity because the salinity is too high or variable for freshwater species or too dilute for most marine species. Thus, most estuarine species have physiological adaptations to cope with the salinity problem. Depending on the hydrography of the estuary, salinity can also vary considerably on a seasonal or even a tidal basis, and variation can be even more limiting to species diversity than mean salinity. Although some estuarine bottoms are rocky, most are at least partially covered by sediments transported from land by rivers and runoff. As is the case with most sedimentary bottoms, the majority of species and individuals live among the sediment grains, concentrated in the upper few centimeters below the sediment–water interface where oxygen is available. Further down in the sediment, oxygen is typically reduced or absent, so that organisms living deeper in the sediment (to tens of centimeters) must be able to tolerate low oxygen or use a long siphon or burrow to maintain oxygen exchange with the surface. Thus, diversity at depth is usually reduced relative to that near the sediment–water interface. Many estuaries are very productive, in part because they are often relatively nutrient rich as a result of the freshwater inflow and terrestrial runoff. As a result, these estuaries often support a very high biomass of a relatively species-poor community. Some coastal habitats are vegetated with photosynthetic plants. Seagrasses are true grasses with roots that occur in shallow subtidal areas, and are


142


usually most abundant in estuaries. Globally, there are only 45 species, but seagrasses provide habitat for other species; these species may include epibionts that live on the seagrass blades, or a variety of fish and sedimentary invertebrates that live below or between the grass blades. Seagrass beds are more diverse, for example, than adjacent sandy bottoms, likely because they provide a predator refuge that a sand bottom cannot. Another common form of vegetation along the seashore is seaweeds. Seaweeds, like seagrasses, require light for photosynthesis and are typically attached to a shallow hard bottom; they are therefore confined to nearshore environments. Some seaweeds are intertidal, showing strong zonation patterns with tidal exposure. Species that are tolerant to prolonged exposure have adaptations that allow them to tolerate the harsh conditions, but again the harsh conditions mean that species numbers in a given intertidal area are relatively few. Kelps make up another dense vegetation habitat that occurs subtidally in cold, clear, nutrient-rich water less than B30m depth. The kelps are attached to the hard bottom, but in some cases sediments may accumulate in areas around the holdfast that attaches the kelp to the substrate. As is the case with seagrasses, a kelp bed is typically dominated by one or two kelp species, but provides critical habitat for many other species. Another feature common to kelps and seagrasses is that they are extremely productive habitats and often support a high biomass of primary and secondary producers. The most spectacular marine habitat in terms of diversity of organisms is the coral reef. Coral reefs are formed by hard corals, which are limited to areas with average surface temperatures above B201C. Reefs offer a wide variety of three-dimensional habitat that is utilized by a variety of other species, resulting in the most diverse marine habitat in terms of number of species per unit area. Many of the most productive marine environments are low in diversity, but reefs are an interesting exception. They are very productive (tight nutrient cycling) as a result of photosynthesizing dinoflagellates called zooxanthellae that are symbionts with reef-building corals. Thus, coral reefs support a high biomass of organisms from corals to fishes. Beyond the immediate near-shore environment lies the continental shelf, which is largely sediment-covered and extends to approximately 200 m depth. The continental shelf varies in width from tens of kilometers to B200 km. Depending on the current and wave regime, sediments may be sandy or muddy, and the sediment composition will influence species composition. Most of the primary production in the

shelf environment is provided by phytoplankton in near-surface waters, and light penetration is typically reduced relative to the open ocean because of increased turbidity and phytoplankton abundance. Benthic vegetation is lacking, but benthic algal diatom mats can sometimes be important where light penetration is sufficient to support them. These mats provide a potential food source for benthic organisms, but contribute little in the form of structural habitat. Epifaunal organisms (those living upon rather than within the seafloor) can provide a habitat that can be important for some species. The productivity of shelf environments is extremely variable geographically, but most of the world’s most-productive fisheries occur on the shelf, particularly where nutrient-rich deep waters are upwelled to the surface. Relative to the near-shore, shelf habitats are usually more diverse, but the level of diversity on shelf environments can vary considerably with geography. For groups as varied as shallow-water molluscs, hard substrate fauna, and deep-sea macrofauna, it has been suggested that species number generally declines from the tropics to the poles. This gradient is much more obvious in the northern hemisphere, where relatively recent glaciation-related extinctions contribute greatly to the pattern. The latitudinal gradient has also been described in the southern hemisphere, though the pattern is less striking. Not surprisingly, given their relative ages, the older Antarctic benthos is more diverse than its Arctic counterpart. But a simple latitudinal gradient is not evident in all taxa. Some groups, such as macroalgae, appear to be most diverse at temperate latitudes, and emerging evidence suggests that shallow-water and deep-sea nematodes may also peak in diversity at temperate latitudes.

The Open Ocean Beyond the continental shelf lies the open ocean, or the pelagic environment. Pelagic communities are delineated primarily by water masses, so that assemblages are often broadly distributed over ranges that may be characterized by differences in temperature, salinity, and nutrient values. Thus, surface circulation and wind effects can interact with thermohaline circulation patterns to create distinct water masses with distinct faunas. Interesting processes occur where these water masses meet, creating complex spatial patterns at the boundaries. Clearly these environmental variables play a major role in regulating pattern, but the geological history of a habitat can also play a role by determining the



regional species pool. Major geological events such as the elevation of the Panama Isthmus, the opening of the Drake Passage, and Pleistocene glaciation all had profound effects on circulation, which in turn had major effects on marine distribution patterns that are reflected in modern communities. Unlike shelf and coastal environments, the influence of benthic communities on water column processes in the open ocean is minimal and indirect. Vertically migrating species play a significant role in oceanic pelagic communities and provide a conduit between surface and deep waters. Some species migrate many hundreds of meters on a daily basis, thus complicating efforts to evaluate biodiversity in a given water mass. These migrating species provide a means of energy transfer between waters at different depths and also provide a mechanism by which regulation of diversity pattern in surface waters could be related to diversity patterns in deeper waters or the benthos. Many benthic species also produce larval stages that may contribute to the biodiversity of surface waters, often on a seasonal basis. Surface waters may also have a major impact on benthic communities in the deep sea because the deep sea depends largely on surface primary production. Photosynthesis is limited to the upper portion of the water column (B200 m) where there is sufficient light penetration. Beneath these waters are the continental slope (200–3000 m), the continental rise (3000–4000 m), the abyssal plains (4000–6500 m), and trenches (6500–10 000 m) of the deep ocean, which will be grouped here as ‘deep-sea’ habitats. Like the shelf environment, most of the deep-sea bottom is covered in sediments, some of which are geologically derived and others that have formed from sinking skeletons of pelagic organisms. As described earlier, pelagic communities are delineated primarily by water masses, and a number of biogeographic provinces have been described for the world’s oceans. Shelf and offshore communities are markedly different in composition and abundance. Local shelf communities are less species rich than offshore communities, but greater spatial heterogeneity in near-shore environments typically results in greater total species richness in neritic environments. One major variable that affects diversity is productivity; species richness tends to be depressed in areas where productivity is high and seasonally variable. This pattern may explain the general pattern of decreasing species number with increased latitude. Characteristics of the regional circulation can offset this general pattern, particularly in coastal environments where productivity does not show as clear a relationship with latitude. Variation in

143

diversity has also been observed with depth in the water column, which is also consistent with primary productivity being restricted to the lighted surface layer of the ocean. The shallowest areas of the oligotrophic North Pacific are more productive in terms of phytoplankton than deeper waters, and phytoplankton diversity is higher in deeper waters than near the surface. Zooplankton, by contrast, show a slightly different pattern where species numbers are higher in surface waters. Studies from the North Atlantic suggest that species richness increases and then peaks at approximately 1000 m. One critical variable that may contribute to these differences in pattern is the pulsed nature of organic input in some systems; highly variable organic flux may represent a strong disturbance, and therefore depress diversity. Because water depth is so great, much of the water column and benthic environment is devoid of light, and the food source is the material sinking from surface waters and advected from the adjacent shelf habitat. The great depths also result in ambient pressures far in excess of those in shallow water, and water temperatures are relatively low (o41C) and are seasonally and spatially much less variant than in shallow-water systems. The seeming inhospitable nature of the deep-sea environment led some earlier investigators to speculate that it was azoic, or devoid of life. Work by Hessler and Sanders in the 1960s and more recent work by Grassle and Maciolek in the 1980s has dramatically changed this view. Although the densities of organisms that live in deepsea sediments are very low and individuals tend to be very small, the numbers of species present is usually very high. Thus, a given sample will contain few individuals, but many of them will represent different species. This generalization is true for most deepsea habitats, but areas such as trenches, upwelling areas, areas with intense currents, and high latitudes can be low in diversity. Low diversity in these areas results from some overwhelming environmental variable, such as low oxygen, resulting in the exclusion of many species. Depth-related patterns have also been described in benthic communities. A peak in biodiversity has been described at continental slope depths, with lower diversity at shelf and abyssal plain depths. This pattern depends on how diversity is defined. In terms of total numbers of species per unit area, shallow water habitats sometimes have higher values because they support much higher densities of individuals. Shallow-water habitats are also patchier over spatial scales of tens to hundreds of kilometers in terms of sediment type, and other habitat variables that change species composition. Thus,


144


pooling of samples from coastal environments can sometimes produce greater total species numbers than pooling over similar distances in deep-sea sediments. For most areas that have been sampled, there is another key difference between deep-sea and shallow-water samples. An individual deep-sea sample is typically characterized by low dominance and greater dissimilarity between proximate samples than would be found in shallow-water samples. One other key difference is total area; although densities of organisms in the deep sea are much lower than in shallow water, the tremendous area of the deep sea alone is enough to support very large numbers of species. But some recent evidence from Australian shelf samples suggests that some shallow-water communities may rival deep-sea communities even at the sample scale. Thus, poorly sampled areas such as tropical shorelines and the southern hemisphere need to be better understood before definitive ‘rules’ on biodiversity patterns can be established. The one major exception to the generality of low productivity in the deep sea is hydrothermal vents, which were first discovered in 1977. Their discovery came as a great surprise to deep-sea scientists because they supported high abundances of novel megafaunal species. The size and numbers of vent organisms is in sharp contrast to most deep-sea habitats, and is possible only because of the chemosynthetic bacteria that form mats or live symbiotically with several vent species. These bacteria are dependent on hydrogen sulfide and other reduced compounds emitted at vents. Since the discovery of vents over 20 new families, 100 new genera, and 200 new species have been described from these communities, but diversity is very low as a result of the toxic hydrogen sulfide. Vent habitats have extremely high levels of endemism resulting from the evolution of forms that are able to thrive in the toxic conditions and take advantage of the high levels of bacterial chemosynthetic production that drives the food chain at vents. Moreover, the fauna at vents is very distinct from other habitats, sometimes at the family level or higher. Among deep-sea habitats, there are other low-diversity communities. Low diversity is also observed beneath upwelling regions, where high levels of organic matter sinking from surface water to bottom sediments may create hypoxic conditions that eliminate many species. Deep-sea trenches are subject to slumping events that contribute to relatively low numbers of species. Deep-sea areas in the Arctic are also still rebounding from loss of much of the fauna during glaciation and anoxic periods that were associated with that time period.

Regulation of Pattern and Linkages Between Organisms and Habitats One pattern common to marine communities and terrestrial communities alike is that habitats characterized by high levels of disturbance, or conditions that are extremely challenging from a physiological perspective, often support relatively few species. Aside from habitats that are strongly influenced by overriding environmental variables that depress diversity, regulation of diversity in marine ecosystems is not fully understood. As a general rule in ecology, high habitat complexity is thought to support the highest diversity because of the many available habitats and niches. Certainly the high level of habitat complexity observed in reefs contributes to their diversity, but pelagic habitats and deep-sea sediments would appear at first glance to be among the least spatially complex habitats in nature. But in these environments, complexity may occur at small temporal and spatial scales. In the pelagic realm, microstructure of nutrients is thought to be an important aspect of species coexistence. Thus, where nutrient levels are relatively low and patchy, primary producers are most diverse. Patchiness is also thought to be important in maintaining diversity in coral reefs and deep-sea communities, the two most speciose community types in the oceans. In both cases, it is thought that intermediate levels of disturbance may be important in preventing the strongest competitors from taking over and eliminating the weaker competitors. In reef habitats, periodic small-scale disturbance in the form of storms keeps any one species from taking over. In the deep sea, it is thought that small-scale disturbances in the form of food patches, biological structures, and sediment topography, can all create microhabitats that allow species to coexist. Determining the factors that regulate biodiversity in marine systems is, nonetheless, an ongoing research question.

Estimates of Total Species Numbers Because such a large portion of the ocean is poorly sampled, estimates of total species numbers vary considerably depending on the specific assumptions used to extrapolate from those areas that have been well sampled. The number of described species from the oceans is approximately 300 000. For some areas of the ocean, such as coastal environments of the North Atlantic, most faunal groups are reasonably well described. However, major gaps in our knowledge exist for some taxonomic groups and some marine habitats. In terms of taxonomic groups, microbes are very poorly known but new evidence



based on molecular approaches suggests that the pelagic and sedimentary realms may both support numbers of bacterial species that would add greatly to the approximately 1.75 million species of nonmicrobial species presently described globally. Protistan diversity is also poorly known. Surprisingly, even for relatively well-sampled areas of the oceans, diversity in some groups of organisms, such as the nematodes, remains poorly known. The specific habitats that are very poorly known include tropical sediments, coral reefs, and deep-sea sediments. Reaka-Kudla has estimated that up to 9 million species may inhabit coral reefs, and Grassle and Maciolek estimated that there may be 10 million macrofaunal species in the deep sea. Others have extrapolated from the ratio of known to unknown species in specific areas to generate estimates of 500 000 to 5 million macrofaunal species in deep-sea sediments alone. Based on the typical ratio of nematode species to macrofaunal species, Lambshead hypothesized that there may be as many as 100 million nematode species in the deep sea. The problem with these estimates is that they are based on a relatively small area, so that the error in extrapolating to the whole of the deep sea, or all oceans, is quite large. For example, it has been estimated that deep-sea sediments (defined here as shelf edge and greater depths) cover approximately 65% of the Earth’s surface yet globally only B2 km2 of ocean floor has been sampled for macrofauna, and B5 m2 has been sampled for meiofauna.

Threats Different habitats presently face different levels of threat as a result of human activity. Open ocean habitats, both pelagic and benthic, have been least impacted by human activity. The greater distance from human populations and their influences, and the shear size of the habitats themselves, make them much less vulnerable than the shoreline and shelf habitat where most marine habitat destruction and local species loss is occurring. Most of the world’s fisheries are concentrated in shelf or near-shore environments, although the capacity to fish deeper waters is increasing all the time. It is estimated that >65% of global fisheries are either fully or overexploited. There are a variety of mechanisms by which this fishing activity may affect biodiversity of marine systems. One of the greatest concerns in terms of benthic habitats is the habitat destruction caused by fishing gear that is dragged across the sea floor, damaging organisms that live in and upon the seabed, homogenizing bottom habitats,

145

and disrupting the sediment fabric and its geochemistry and microhabitats. Recent estimates suggest that some major fishing grounds, such as Georges Bank, have experienced trawl coverage that exceeds 200–400% annually. Habitat damage is not an issue for the fluid pelagic habitat, but the removal of nontarget species as by-catch remains a problem in bottom and pelagic fisheries. Organisms ranging from sea birds to marine mammals to fish to invertebrates are all known to suffer high by-catch mortality; by-catch levels can often rival or even exceed the biomass of the target species removed from an ecosystem. Another major concern with fisheries is that they are often very effective at removing large numbers of individuals of the target species, which is often a top predator within the ecosystem. The potential for alteration of food chains is great, both for pelagic and benthic species. Evidence is accumulating that the trophic structure of many heavily fished ecosystems is changing, with upper trophic levels sometimes being eliminated and the transfer of energy through remaining species altered substantially. One final aspect of fishing activity that is of particular concern for pelagic habitats is lost fishing gear and debris that may continue to ‘ghost fish’ and capture target and nontarget organisms for years after the gear has been lost. One chronic problem with studies of fisheries impacts is that we lack good ‘control’ sites where fishing impacts are minimal. Advances in fishing technology have made most habitats accessible to fishing gear, and the few areas that remain inaccessible are typically poor ‘control’ sites because they are fundamentally different in physical topography and species composition than the areas that are fished. Coral reef fisheries have their own unique problems. Dynamite, cyanide, and bleach fishing are all used to get around the problem of fishing a topographically complex habitat, but they are also methods that destroy many nontarget species including the corals that make up the habitat itself. One of the discouraging facts about reef systems is that in most instances, reefs were considerably altered by removal of large and potentially important species by early settlers, and we have little idea of what the pristine systems looked like in terms of species relationships. Aquaculture represents a special case in terms of fishing activity, in that impacts are typically more localized and easily seen. The issue of ownership is also less contentious, stocks are more easily managed, and in some cases it is also easier to assign accountability for environmental damage. But aquaculture can be extremely destructive, in that some activities such as shrimp farming are often


146


achieved by destroying other coastal habitats such as mangroves. Aquaculture also often involves moving organisms around as brood stock, potentially allowing invasion of nonnative habitat by the species being cultured, or parasites and diseases that the organism may carry into the new environment. Even for those organisms that have been deliberately moved, there are potential consequences in terms of alteration of local genetic structure of natural populations. Many coastal environments are also threatened by pollution resulting from the dumping of various toxic wastes such as heavy metals, organic compounds (e.g., polychlorinated biphenyls or PCBs), metals, and eutrophication resulting from increased input of macronutrients from sewage and agricultural runoff. These excess nutrients result in blooms of a low diversity phytoplankton community (sometimes favoring toxic species) that sink to the bottom and undergo microbial decomposition. In some instances, microbial respiration will deplete bottom waters of oxygen, and benthic communities may subsequently be wiped out. Whether the cause is eutrophication or toxic chemicals, both forms of pollution depress diversity and favor a few weedy species. On coral reefs, increased nutrients will typically lead to macroalgal blooms and the loss of corals. In this instance, the entire habitat may be destroyed. The effect of fisheries operations on habitat alteration has already been mentioned, but marine habitats may also be altered for other activities. The demand for coastal real estate, both for industrial and residential use, has eliminated large areas of salt marsh and mangrove. Expansion of coastal waterways by dredging and replenishing of eroded beaches by mining subtidal sands are two mechanisms by which habitats may be altered. Invariably, the loss of habitat means the loss of species, at least on a local scale. Unfortunately, the amount of habitat being lost is becoming so extensive that the number of habitat refugia remaining are becoming fewer and fewer and more widely separated. There has been great concern over the introduction of non-native (‘exotic’) species into marine habitats in recent years, a phenomenon that has been ongoing for some time. Indeed, it has been estimated that between the years 1500 and 1800, more than 1000 intertidal and subtidal species worldwide may have been transported and introduced into nonnative habitats. As many as 3000 species may be in transit in the ballast water of ships on a given day. In some instances it has been documented that invasive species have greatly altered the species composition and ecological

processes in the areas they have invaded. San Francisco Bay, for example, now harbors hundreds of exotic species, including Asian clams that have become so abundant that they have altered the seasonal phytoplankton production cycle within the bay. The problem of invasive species is thought to be most severe in coastal and estuarine regions, where open ocean waters have historically provided a barrier to dispersal. Ballast water is not, however, the only culprit in facilitating the movement of invasive species. Fouling organisms on the hulls of ships provide another mechanism, and aquaculture and scientific study are additional transport vectors. One other threat to marine biodiversity is global climate change. The effects of climate change are difficult to know given the different trajectories predicted by different models, but several general categories of effect may be expected. First, as air temperatures increase, so will ocean surface temperatures, presumably shifting faunas toward the poles. Such an effect has already been documented in California intertidal and pelagic communities. What is less obvious is that some species will be unable to simply shift poleward because other habitat requirements, such as the presence of a shallow bank, may not be met at another latitude. Coral reefs provide an excellent example of this phenomenon. Recently, large areas of coral reefs have experienced increased occurrences of bleaching, a phenomenon where corals expel their symbiotic dinoflagellates and die. Bleaching has been linked to the increased frequency of El Ninõ events, higher water temperature, and elevated ultraviolet radiation; in this case, corals cannot simply colonize an adjacent habitat because it is typically too deep and warming trends are much faster than potential colonization rates. A second category of effect is rising sea level. In some instances it may be possible that intertidal organisms may simply shift upward in response to rising waters, but in areas where human populations have developed areas in the landward direction, mangroves, salt marshes, and intertidal habitats may be prevented from advancing by seawalls or other physical barriers. Perhaps the most difficult aspect of global change to predict is the effect that alteration of temperature and rainfall pattern will have on ocean circulation. Some global warming models have suggested that even relatively modest temperature increases may alter ocean circulation patterns. Because ocean circulation is a critical variable in terms of surface productivity and transport of reproductive propagules, any change to ocean circulation could potentially affect every marine community from shallow surface waters to the deepest areas of the ocean.



147

The Importance of Marine Biodiversity and clear water. Other aesthetic services are provided Marine organisms contribute to many critical processes that have direct and indirect effects on the health of the oceans and humans. In the majority of instances there are few data to demonstrate that total numbers of species are important, but data on this question are only starting to be assembled. What is obvious is that there are specific species and functional groups that play very critical roles in important ecosystem processes, and the loss of these species may have significant repercussions for the whole ecosystem. Primary and secondary production are critical mechanisms by which marine communities contribute to global processes. It has been estimated that approximately half the primary production on Earth is attributable to marine organisms. It has also been estimated that approximately 20% of animal protein consumed by humans is from the oceans. Perhaps more importantly, this consumption is much higher in some countries than in others, making it a critical staple of many diets. Marine organisms are also harvested for extractable products, including medicines and various industrial products. The global cycling of nutrients and even carbon depend on marine communities. Without primary producers in surface waters, the oceans would quickly run out of food, but without planktonic and benthic organisms to facilitate nutrient cycling, the primary producers would quickly become nutrient limited. Benthic marine organisms can contribute to sediment and shoreline stability. Mangroves, salt marsh plants, and seagrasses all bind sediments together and thereby reduce shoreline erosion. Sedimentary organisms can either stabilize or destabilize sediments, thus affecting coastal sediment budgets. One direct ramification of these activities is that sediment bound pollutants will be greatly affected by binding and destabilization of sediments, creating a situation where sedimentary bacteria, diatoms, and invertebrates can influence whether pollutants are buried or resuspended. Some bacteria and invertebrates are also able to metabolize certain pollutants and reduce or eliminate their toxicity. In coastal environments, salt marshes, mangroves, and seagrasses can help trap sediments and absorb nutrients from sewage and agricultural sources, thereby filtering coastal runoff and helping to maintain relatively non-turbid, non-eutrophied coastal waters. In a related manner, benthic organisms may be important filter feeders, removing particles that would reduce water clarity. In addition to allowing light penetration to greater depths, this filtering activity can contribute to coastal aesthetics

by coastal wetlands and coral reefs, both of which generate tremendous tourist interest.

Concluding Remarks Although documented extinctions are relatively few, there are good reasons for improving our understanding of biodiversity pattern and regulation, and the role that biodiversity plays in key ecosystem processes. One problem is that biodiversity in the oceans is poorly described, and the handful of seabirds, marine mammals, and marine snails that have become extinct may represent the visible few of a much larger number of organisms that have been lost without our knowing it. It has been estimated, for example, that 50 000 undescribed species may have already been eliminated from coral reefs. A second concern is that some of the species we have treated as cosmopolitan may, in fact, represent a sibling species group, and when we eliminate a species from an area, it may be a different species than occurs elsewhere. A third concern is that even if the elimination of species from a given area does not represent a global extinction, it may represent a unique genotype that cannot be replaced from a surviving population elsewhere. Although several general patterns of biodiversity have been described in the oceans, and our understanding of how biodiversity is regulated and maintained is still limited, several paradigms point to the importance of disturbance and habitat heterogeneity. Because vast areas of the oceans remain unsampled, our estimates of species numbers are crude and based on a very small portion of the marine habitat, but they do suggest that the oceans contain a significant portion of the global species pool. Unfortunately, we know little about how biodiversity contributes to the many critical ecosystem services that marine communities provide. In many instances, it is possible to generate very clear examples of a single species having a major impact on its ecosystem, and based on this observation it is safe to say that loss of biodiversity may have very negative effects on marine ecosystems if the species lost is one of these key players. But we have little idea of whether numbers of species matter, or which species are most important in maintaining a properly functioning ecosystem. An improved understanding of biodiversity and regulation will give us better tools to predict where biodiversity changes are likely to occur and why, how these changes will affect other components of marine biodiversity, and the best strategy to mitigate such changes.


148


See also Benthic Boundary Layer Effects. Benthic Foraminifera. Benthic Organisms Overview. Deep-Sea Fauna. Fish: Demersal Fish (Life Histories, Behavior, Adaptations). Macrobenthos. Meiobenthos. Ocean Margin Sediments.

Further Reading Angel MV (1997) Pelagic biodiversity. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity. Patterns and Processes. Cambridge: Cambridge University Press. Birkeland C (ed.) (1997) Life and Death of Coral Reefs. New York: Chapman and Hall. Costanza R, d’Arge R, de Groot R, et al. (1997) The value of the world’s ecosystem services and natural capital. Nature 387: 253--260. Dayton PK, Thrush SF, Agardy MT, and Hofman RJ (1995) Environmental effects of marine fishing. Aquatic Conservation: Marine and Freshwater Ecosystems 5: 205--232. Gage JD and Tyler PA (1991) Deep-Sea Biology: A Natural History of Organisms at the Deep-Sea Floor. Cambridge: Cambridge University Press. Grassle JF, Lasserre P, McIntyre AD, and Ray GC (1991) Marine biodiversity and ecosystem function: a proposal for an international programme of research. Biology International Special Issue 23: 1--19. Grassle JF and Maciolek NJ (1992) Deep-sea species richness: regional and local diversity estimates from

quantitative bottom samples. American Naturalist 139: 313--341. Gray JS (1997) Marine biodiversity: patterns, threats and conservation needs. Biodiversity and Conservation 6: 153--175. Gray J, Poore G, and Ugland K (1997) Coastal and deepsea benthic diversities compared. Marine Ecology Progress Series 159: 97--103. Hall SJ (1999) The Effects of Fishing on Ecosystems and Communities. Oxford: Blackwell Science. Lambshead PJD (1993) Recent developments in marine benthic biodiversity research. Oceanis 19: 5--24. May R (1992) Bottoms up for the oceans. Nature 357: 278--279. McGowan JA and Walker PW (1985) Dominance and diversity maintenance in an oceanic ecosystem. Ecological Monographs 55: 103--118. National Research Council (1995) Understanding Marine Biodiversity: a Research Agenda for the Nation. Washington, DC: National Academy Press. Norse EA (ed.) (1993) Global Marine Biological Diversity: a Strategy for Building Conservation into Decision Making. Washington, DC: Island Press. Rex MA, Etter RJ, and Stuart CT (1997) Large-scale patterns of biodiversity in the deep-sea benthos. In: Ormond RFG, Gage JD, and Angel MV (eds.) Marine Biodiversity. Patterns and Processes. Cambridge: Cambridge University Press. Snelgrove PVR, Blackburn TH, Hutchings PA, et al. (1997) The importance of marine sedimentary biodiversity in ecosystem processes. Ambio 26: 578--583.


DOLPHINS AND PORPOISES R. S. Wells, Chicago Zoological Society, Sarasota, FL, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction: Physical Descriptions and Systematics Dolphins and porpoises are members of a diverse group of cetaceans including 44 species belonging to 6 of the 10 families of modern toothed whales, of the suborder Odontoceti (Table 1). In general, this group includes the smaller odontocetes, with body lengths ranging across species from about 1.5 m up to about 9 m. The grouping represents two major and very different subgroups. Dolphins include 35 species of the families Delphinidae and Pontoporiidae, primarily inhabiting inshore and/or pelagic marine waters worldwide, and three families of obligate river dolphins (Iniidae, Lipotidae, and Platanistidae), each represented by a single species. The distributions of the three river dolphin species are limited to riverine systems in Asia or South America. Porpoises include the six species of the family Phocoenidae, also primarily inhabiting inshore and pelagic waters. Dolphins and porpoises exhibit essentially the same basic streamlined spindle or fusiform body shape, which provides for effective movement through the dense water medium (Figure 1). They have a single blowhole on top of their head for respiration. In their skulls, they both demonstrate telescoping overlap of the maxillary bones over the frontals in the supraorbital region. Most dolphins and porpoises have numerous teeth (polydonty), and these teeth are homodont (all alike in structure) in all but one species. A few small hairs occur on the beaks of the animals at the time of birth, but these are soon lost. Postcranially, the bones of the hand and arm have been modified into a pectoral flipper, a solid winglike structure articulating with the shoulder, and serving as a control surface for maneuvering. Pelvic appendages have been reduced to small, internal pelvic bones (although a bottlenose dolphin captured in Japan in 2006 presented a bilateral pair of fully formed, small fins in the pelvic region). The vertebral column, with its variably fused cervical vertebrae and prominent processes for attachment of the strong musculature used for up and down movement of the tail for propulsion, tapers along the tail stock

or peduncle, toward the tail to the flukes, a pair of fibrous horizontal fins. A fibrous dorsal fin located near the middle of the back of most, but not all, dolphins and porpoises serves to stabilize the animals as they swim, as a radiator for cooling the reproductive organs, and as a weapon. This structure can also be useful to researchers, as the dorsal fins of many cetaceans have individually distinctive shapes and natural notch patterns, providing a means of reliable identification for recognizing individuals through time. Males and females look quite similar and are of similar size in most dolphin and porpoise species, though there may be differences in body and/or appendage size in some, especially among the larger dolphins. Female dolphins typically have a genital and an anal opening in a single ventral groove, with a nipple located in a mammary slit on each side of the genital opening. Males typically have a genital opening in a ventral groove anterior to the separate groove for the anus. Where it occurs, sexual dimorphism is variable across the species, with females of some of the smaller species (Cephalorhynchus spp., the tucuxi, Sotalia fluviatilis) being slightly larger than the males, whereas the males of the largest species (killer whales, false killer whales (Pseudorca crassidens), pilot whales (Globicephala spp.)) tend to be much larger than the females. In some cases, features that might be used in competitive displays or affect performance in battle or chasing potential mates, such as dorsal fin height (e.g., killer whales, Figure 2), peduncle height, or fluke span, are disproportionately larger for males than for females. Smaller species such as spinner dolphins can demonstrate some degree of dimorphism, with adult males having taller and pointier dorsal fins, and a postanal keel or hump on the ventrum. Most dolphins and porpoises exhibit countershading, with a darker dorsum and lighter ventrum. Countershading presumably functions as camouflage to facilitate approaching prey and avoiding predators. Dolphins

There is much variation among the species of dolphins regarding robustness, the presence of a clearly demarcated projecting beak, and the length of the beak, numbers and size of teeth, and the presence of a dorsal fin. The characteristic cone-shaped, or peglike teeth of dolphins differ across the species in size


149


18 19 20 21 22 23

10 11 12 13 14 15 16 17

7 8 9

6

2 3 4 5

1

Sp #

Table 1

L. obscurus subsp. Lissodelphis borealis Lissodelphins peronii Orcaella brevirostris Orcinus orca Peponocephala electra Pseudorca crassidens

L. obscurus obscurus

L. obscurus fitzroyi

Grampus griseus Lagenodelphis hosei Lagenorhynchus acutus Lagenorhynchus albirostris Lagenorhynchus australis Lagenorhynchus cruciger Lagenorhynchus obliquidens Lagenorhynchus obscurus

G. melas melas G. melas subsp. G. melas edwardii

Feresa attenuata Globicephala macrorynchus Globicephala melas

Falklands, S. American dusky dolphin S. African, Indian Ocean dusky dolphin New Zealand dusky dolphin N. right whale dolphin S. right whale dolphin Irawaddy dolphin Killer whale, or orca Melon-headed whale False killer whale

N. Atlantic long-finned pilot whale N. Pacific long-finned pilot whale S. Hemisphere long-finned pilot whale Risso’s dolphin, or Grampus Fraser’s dolphin Atlantic white-sided dolphin White-beaked dolphin Peale’s dolphin Hourglass dolphin Pacific white-sided dolphin Dusky dolphin

Pygmy killer whale Short-finned pilot whale Long-finned pilot whale

Short-beaked common dolphin

Marine Dolphins Commerson’s dolphin Falklands, S. American subspecies Kerguelen subspecies Chilean dolphin Heaviside’s dolphin Hector’s dolphin Long-beaked common dolphin

Family Delphinidae Cephalorhynchus commersonii C. commersonii commersonii C. commersonii subsp. Cephalorhynchus eutropia Cephalorhynchus heavisidii Cephalorhynchus hectori Delphinus capensis

Delphinus delphis

Common name

Taxon

List of extant species and subspecies of dolphins and porpoises

New Zealand N. Pacific Ocean S. Pacific, S. Atlantic, S. Indian, and Southern Oceans Coastal Indo-Pacific waters Worldwide Worldwide, tropical to temperate waters Worldwide, tropical through temperate waters

S. Africa

Worldwide, tropical through temperate waters Tropical and warm temperate waters of all oceans Colder waters of the N. Atlantic Colder waters of the N. Atlantic Southern S. America, S. Atlantic and S. Pacific Oceans Worldwide, Southern Ocean N. Pacific Ocean Southern S. America, Atlantic and Pacific Oceans, S. Africa, New Zealand Southern S. America, Falkland Islands

Southeastern S. America, Kerguelen and Falkland Islands Southeastern S. America, Falkland Islands Kerguelen Island Southwestern S. America Southwestern Africa New Zealand Tropical and warm temperate waters of the Pacific, Indian, and S. Atlantic Oceans Tropical to temperate waters of the Pacific and N. Atlantic Oceans Worldwide tropical and warm temperate waters Worldwide tropical to temperate waters N. Atlantic and southern S. Pacific, S. Indian, and S. Atlantic Oceans N. Atlantic Ocean N. Pacific Ocean S. hemisphere

General range

NE LC DD DD LR(cd) LC LC

NE

NE

DD DD LC LC DD LC LC DD

NE NE (prob. extinct) NE

DD LR(cd) LC

LC

DD NE NE DD DD EN LC

IUCN designationa

150 DOLPHINS AND PORPOISES

Stenella coeruleoalba

Stenella frontalis

Stenella longirostris S. longirstris longirostris S. longirostris orientalis S. longirostris centroamericana

29

30

31

(c) 2011 Elsevier Inc. All Rights Reserved. La Plata Dolphin Franciscana South American River Dolphins Amazon dolphin, or boto Orinoco dolphin Amazon dolphin Bolivian dolphin Chinese River Dolphin Baiji or Yangtze dolphin S. Asian River Dolphins Susu, or ‘blind’ river dolphin Ganges dolphin Indus dolphin

Tursiops aduncus

Family Pontoporiidae Pontoporia blainvillei

Family Iniidae Inia geoffrensis I. geoffrensis humboldtiana I. geoffrensis geoffrensis I. geoffrensis boliviensis

Family Lipotidae Lipotes vexillifer

Family Platanistidae Platanista gangetica P. gangetica gangetica P. gangetica minor

34

1

1

1

1

Indo-Pacific bottlenose dolphin

Tursiops truncatus

33

Common bottlenose dolphin

S. longirostris roseiventiris Steno bredanensis

Spinner dolphin Gray’s spinner dolphin Eastern spinner dolphin Costa Rican, Central American spinner dolphin Dwarf spinner dolphin Rough-toothed dolphin, or Steno

Atlantic spotted dolphin

Striped dolphin

Tucuxi Marine tucuxi Freshwater tucuxi Indo-Pacific hump-backed dolphin Atlantic hump-backed dolphin Pantropical spotted dolphin E. Pacific offshore spotted dolphin Hawaiian spotted dolphin E. Pacific coastal spotted dolphin Clymene dolphin

32

28

25 26 27

Sotalia fluviatilis S. fluviatilis guianensis S. fluviatilis fluviatilis Sousa chinensis Sousa teuszi Stenella attenuata S. attenuata subspecies A S. attenuata subspecies B S. attenuata graffmani Stenella clymene

24

Rivers of India, Pakistan, Nepal, Bangladesh Ganges–Brahmaputra River system Indus River

Yangtze River

S. American rivers Orinoco River basin Amazon River basin Upper Rio Madeira drainage

Coastal waters of Argentina, Brazil, Uruguay

Southeast Asian and N. Australian shallow waters Tropical to temperate waters of all oceans, including the Mediterranean Sea Tropical to temperate waters of all oceans and seas, especially near the coast Tropical to temperate coastal waters of the Indian Ocean

Northeastern S. America Coastal marine waters S. American rivers Indian Ocean and Indo-Pacific coastal waters Northwestern African coastal waters Tropical and warm temperate waters of all oceans Eastern tropical Pacific Ocean Hawaiian Islands E. Pacific Ocean, coastal off Mexico to Colombia Tropical and warm temperate Atlantic Ocean and Gulf of Mexico Tropical and warm temperate waters worldwide, including the Mediterranean Sea Tropical to temperate waters of the Atlantic Ocean and Gulf of Mexico Tropical and warm waters of all oceans Tropical and warm waters of all oceans Pelagic waters of eastern Pacific Ocean Pacific Ocean, over continental shelf off Central America

EN EN EN (Continued )

CR (likely extinct)

VU NR NE NE

DD

DD

DD

LC DD

LR(cd) NE NE NE

DD

LR(cd)

DD NE NE DD DD LR(cd) NE NE NE DD

DOLPHINS AND PORPOISES 151

Porpoises Finless porpoise Indian Ocean finless porpoise W. Pacific finless porpoise Yangtze River finless porpoise Harbor porpoise N. Atlantic harbor porpoise W. N. Pacific harbor porpoise E. N. Pacific harbor porpoise Spectacled porpoise Vaquita, or Gulf of California harbor porpoise Burmeister’s porpoise Dall’s porpoise Dalli-phase Dall’s porpoise Truei-phase Dall’s porpoise

Family Phocoenidae Neophocaena phocaenoides N. phocoaenoides phocaenoides N. phocoaenoides sunameri N. phocaenoides asiaeorientalis Phocoena phocoena

P. phocoena phocoena P. phocoena subsp. P. phocoena vomerina Phocoena dioptrica Phocoena sinus

Phocoena spinipinnis Phocoenoides dalli P. dalli dalli P. dalli truei

Common name

Taxon

Continued

Southern S. America, S. Atlantic and S. Pacific Oceans N. Pacific Ocean N. Pacific Ocean Western N. Pacific Ocean

Coastal tropical and warm temperate Indo-pacific waters Coastal waters of the Indian Ocean to the S. China Sea E. China Sea to N. Japan Yangtze River N. Atlantic and N. Pacific Oceans, North, Bering, Barents, and Black Seas N. Atlantic Ocean Western N. Pacific Ocean Eastern N. Pacific Ocean SW Atlantic and Southern Oceans Gulf of California

General range

DD LR(cd) NE NE

NE NE NE DD CR

DD NE NE EN VU

IUCN designationa

Categories include: NE, not evaluated; DD, data deficient; VU, vulnerable; EN, endangered; CR, critically endangered; LR (cd), lower risk, conservation dependent; LC, least concern. Adapted from Reeves RR, Smith BD, Crespo EA, and Notorbartolo di Sciara G (eds.) (2003) Dolphins, Whales, and Porpoises: 2002–2010 Conservation Action Plan for the World’s Cetaceans. Gland: IUCN/SSC Cetacean Specialist Group.

a

5 6

3 4

2

1

Sp #

Table 1

152 DOLPHINS AND PORPOISES


DOLPHINS AND PORPOISES

153

Figure 3 Dentition of a bottlenose dolphin, showing rows of homodont teeth, used to grasp prey. Photo by Randall S. Wells.

Figure 1 The basic body form of dolphins is exemplified by the bottlenose. Photo by Randall S. Wells.

Figure 4 Adult female Franciscana dolphin. Note small body size relative to adult human hand in frame. Photo by Randall S. Wells.

Figure 2 Dorsal fin development in killer whales. Adult males (left) develop much taller and more triangular dorsal fins than subadult males or adult females. Photo by Randall S. Wells.

and number depending on the prey. The pointy teeth are designed for grasping individual prey items, rather than for chewing (Figure 3). The number and size of the teeth in turn influence the size of the beak and shape of the mouth of each species. For example, spinner dolphins (Stenella longirostris) may have more than 200 small teeth in long, pincers-like jaws for capturing small fish and invertebrates associated with the deep scattering layer. In contrast, killer whales (Orcinus orca) have about 50 large teeth for catching large fish and pinnipeds and removing chunks of flesh from a variety of marine mammals. Risso’s dolphins (Grampus griseus) lack a pronounced beak and have only 10 teeth, all in the lower jaw, for grasping soft-bodied squid prey.

The generally falcate shape of the dorsal fin is a characteristic that distinguishes dolphins from porpoises. Within the dolphins, dorsal fins vary from species to species in height and shape, from the 2-2 m tall triangular fin of adult male killer whales, to the absence of dorsal fins in two species of dolphins (Lissodelphis spp.), a rounded fin in Hector’s dolphin (Cephalorhynchus hectori) and the occurrence of a hump at the base of the fin in others (Sousa spp.). Dolphins range in size from the tiny Franciscana (Pontoporia blainvillei) (Figure 4), dwarf spinner (Stenella longirostris roseiventris), and Hector’s dolphins at 1.4–1.5 m and about 50 kg, to adult male killer whales at 9.0 m and 5600 kg. Color patterns of dolphins vary widely from species to species, including patterns of black, white, gray, brown, orange, and/or pink. Water clarity and its effects on the ability for visual communication may contribute to patterns of coloration, as dolphins living in clear, oceanic water tend to have more complex color patterns on their sides than do dolphins living in the murky waters of estuaries or rivers (Figure 5).


154


Figure 5 Common dolphins in the Gulf of California, showing the complex color pattern exhibited by many oceanic dolphins. Photo by Randall S. Wells.

The classification of the numerous members of the family Delphinidae is undergoing much revision with the advent of genetic analysis techniques and the increased efforts by scientists to collect small genetic samples from specimens from around the world. Further confusing taxonomic distinctions for the dolphins, a number of hybrid dolphins (some viable and fertile) have been identified in captive breeding situations and in the wild. Three families/species of peculiar long-snouted dolphins have been linked under the category ‘river dolphin’ because their appearances are somewhat similar to one another, very different from other species of dolphins, and because they spend in their entire lives in rivers. These three species have extremely long beaks, flexible necks, relatively large flukes, and flippers, and their eyes are reduced in size (Figure 6). The term ‘river dolphin’ is perhaps a misnomer for these three single-species families, and it is a hold-over from the past. Though they represent the only obligate freshwater dolphin species, other species (Irrawaddy dolphin, Orcaella brevirostris, and tucuxi, Sotalia fluviatilis) include populations or subspecies that reside permanently in freshwater. Porpoises

As a group, porpoises tend to be the smallest of the cetaceans, with adult body lengths of less than 2.5 m. They have small, rounded braincases and they lack the pronounced rostrum found in many of the dolphin species, contributing to a condition referred to as paedomorphosis, the retention of juvenile characters in the adult form (this is also seen in dolphins of the genus Cephalorhynchus) (Figure 7). Raised protuberances occur on the premaxillae. In contrast to the dolphins, porpoises have spatulate, or spadeshaped teeth.

Figure 6 The Yangtze River dolphin, or baiji, showing the characteristic long beak and reduced eyes. Photo courtesy of Wang Ding, Institute of Hydrobiology, the Chinese Academy of Sciences.

Figure 7 The basic body form of porpoises as exemplified by this vaquita, on the gill net in which it was caught. Photo by Flip Nicklin/Minden Pictures.

Porpoises are stocky and robust, with relatively small appendages. This combination of features may be an adaptation for heat retention for thermoregulation in the cold waters typically inhabited by this family. The dorsal fin, which occurs in all but one (finless porpoise, Neophocaena phocaenoides) of the six species, tends to be low and triangular, as compared to typically falcate dolphin fins. All porpoises except Dall’s porpoise (Phocoenoides dalli) have epidermal tubercles on the leading edge of the dorsal fin. The function of these tubercles is not known. All porpoises have darker patches of pigmentation surrounding the eye, especially the spectacled porpoise (Phocoena dioptrica), but the family generally lacks the broad range of colors found in the dolphins. Earlier confusion regarding the use of the vernacular names dolphin and porpoise seems to have



declined over the past few decades, with increased familiarity with the distinctions between the groups. In some parts of the world, the terms were used interchangeably, especially in reference to coastal bottlenose dolphins. In some cases, members of the family Delphinidae have been referred to as porpoises by fishers and management agencies in order to distinguish the mammals from the dolphin fish, the dorado, or mahi mahi (Coryphaena hippurus).

Evolution The earliest whales, the archaeocetes, appear to have evolved from mesonychian condylarths, ungulate ancestors that were primarily terrestrial. Archaeocete fossils from about 52 to 42 Ma have been found in Africa, India, North America, and Pakistan. The recent fossil discovery of a ‘walking whale’ (Ambulocetus) exemplifies the transition from terrestrial to aquatic life. One of the most distinguishing evolutionary developments from archaeocetes to modern cetaceans was the ‘telescoping’ movement and elongation of the premaxillary and maxillary bones in the skull as the nasal openings migrated to the more effective position on top of the cetacean’s head, creating a rostrum or beak. Odontocetes appear to have diverged from the baleen whales (Mysticetes) about 25–35 Ma. The first modern members of the Delphinidae appeared in the fossil record in the mid–late Miocene, about 10–11 Ma, from kentriodontid-like ancestors, small cetaceans from both the Atlantic and Pacific Oceans that disappeared about 10 Ma. Based on fossil evidence and estimates of divergence between the cytochrome b genes, the families Delphinidae and Phocoenidae appeared at about the same time, and both derived from the Kentrodontidae. The family Pontoporiidae appears to have originated in Pliocene and mid-Miocene marine environments of Peru. The adaptation of the different river dolphins to freshwater environments appears to be a convergence. The Iniidae are believed to have entered the Amazon River basin from the Pacific Ocean 15 Ma, or alternatively from the Atlantic about 1.8–5.0 Ma. Platanistid fossils come from mid-Miocene marine deposits in North America and Europe. The Lipotiidae appear to have originated in the Miocene in the North Pacific, with fossils from China and California.

Life History Little is known about the life history patterns of most dolphins and porpoises. Most of the available information has been derived from examination of

155

carcasses obtained from strandings or fisheries, but some insights have also come from a few long-term studies of wild dolphin populations. Both groups typically produce a single offspring at a time, after a gestation period of about 9–16 months, depending on the species. Larger dolphins, such as pilot whales and killer whales, tend to have the longer gestation periods, with most other species closer to 1 year. In general, members of the family Delphinidae develop at a slower pace than the porpoises. Dolphins tend to mature later in life than porpoises (typically 6–12 years vs. 3 years for Phocoena). At least some porpoises are capable of annual reproduction and can be simultaneously pregnant and lactating, while many dolphins rear calves for multiple years before giving birth to the next calf. Bottlenose dolphins in Sarasota Bay, Florida have been documented to give birth as old as 48 years of age and, with 3–6 -year (successful) calving intervals on average, have been observed with up to eight different calves in a lifetime. It is believed that few porpoises live longer than 20 years, while the maximum lifespan of most dolphins is likely to be at least 20 years and, in the cases of the larger species, may exceed 60 years. The oldest female bottlenose dolphin in Sarasota Bay is estimated to be 56 years old; the oldest male is 48 years old. Females in some of the large, sexually dimorphic species of dolphins may live 15–20 years longer than males, and become postreproductive. Longevity of river dolphins has yet to be determined, but one baiji (Lipotes vexillifer) lived for 22 years in captivity. Mortality and serious injury of dolphins and porpoises result from a variety of natural sources, including pathogens, biotoxins from Harmful Algal Blooms, predation by sharks and killer whales, and stingray barbs. Anthropogenic sources of mortality that have been clearly identified include directed hunts with harpoons, nets, or drive fisheries, incidental mortality in commercial fishing nets and longlines, ingestion of and entanglement in recreational fishing gear, ingestion of foreign objects, and collisions with boats. Though not yet conclusively demonstrated, the weight of evidence suggests that environmental contaminants such as heavy metals (e.g., mercury), organchlorines (e.g., PCBs and DDT pesticides and their metabolites) along with emerging chemicals such as brominated fire retardants may adversely impact health and/or reproduction, and high-amplitude sounds from some military sonars and underwater explosions may kill cetaceans. Dolphin and porpoise life history patterns may change in response to changes in abundance such as those resulting from heavy mortality in commercial fisheries. Density-dependent responses have been


156


observed in a variety of cases, such as for dolphins of the genus Stenella in the tuna seine net fishery in the Eastern Tropical Pacific Ocean (ETP). As dolphin densities have declined, their age at sexual maturity has also declined, presumably in response to more resources becoming available to remaining individuals. A possible extreme case of this involves the Franciscana dolphin of eastern South America, which reaches sexual maturity at less than 3 years of age, the youngest age known for any odontocete cetacean. It has been suggested that this low age at sexual maturity has come about in response to continuing heavy losses in fishing nets.

Distribution, Ranging Patterns, and Habitat Use Dolphins and/or porpoises can be found in nearly every marine habitat in the world, as well as in several major river systems (Table 1). The most cosmopolitan species, the killer whale (Orcinus orca), inhabits waters from the polar ice edge through the tropics. Dolphins reach their highest diversity in tropical and warm temperate waters. Many species are pantropical, while others are limited to one or two ocean basins. There exists only one anti-tropical dolphin species, the long-finned pilot whale (Globicephala melas), but there are several anti-tropical species pairs (such as the northern and southern right whale dolphins, Lissodelphis spp.). Pelagic habitats support a number of dolphin species, as well as Dall’s porpoises (Phocoenoides dalli) and spectacled porpoises (Phocoena dioptrica). Offshore deep water areas with a stable mixed layer and a shallow thermocline are home to members of the genus Stenella and rough-toothed dolphins (Steno bredanensis). More variable offshore areas, where upwelling occurs, is preferred by species such has pilot whales (Globicephala spp.), common dolphins (Delphinus spp.), striped dolphins (Stenella coeruleoalba), and melon-headed whales (Peponocephala electra). Coastal waters, including bays, sounds, and estuaries, are preferred by species such as bottlenose dolphins (Tursiops spp.), hump-backed dolphins (Sousa spp.), Franciscana (Pontoporia blainvillei), harbor porpoises (Phocoena phocoena), vaquita (Phocoena sinus), and Burmeister’s porpoises (Phocoena spinipinnis). Dolphin and porpoise habitats are three dimensional, and physical habitat features as well as prey availability influence habitat use. Diving is an adaptation to life in deeper waters, and diving abilities appear to vary greatly both within and across species of dolphins and porpoises. Few data on

diving capabilities are available. Bottlenose dolphins commonly reside in shallow inshore waters, where deep dives are neither necessary nor possible. They utilize resources associated with seagrass meadows, inlets, and other features. However, where Tursiops occurs offshore, for example, off the island of Bermuda, it has been documented to dive to depths of 1000 m or more. Rehabilitated rough-toothed dolphins (Steno bredanensis) tagged with satellite-linked time-depth recorders and released into pelagic waters rarely dove below 50 m. Based on stomach content data, spinner dolphins (Stenella longirostris) are not thought to dive below about 200–300 m. A rehabilitated and tagged Risso’s dolphin (Grampus griseus) dove occasionally to depths of 400–500 m, but most dives were shallower. None of the porpoises are believed to be deep divers. The true river dolphins are found only in Asia and South America. Three subspecies of Inia spp. inhabit the dynamic Amazon and Orinoco river drainages. These dolphins move from deep channels in time of low water into the rainforest canopy and grasslands during flood season. Two different subspecies of Platanista spp., isolated for at least hundreds of years, inhabit the Ganges and Indus river systems. Historically, the baiji (Lipotes vexillifer) has inhabited the Yangtze River drainage, including associated large lakes and tributaries, but much of this habitat is no longer accessible to the dolphins due to damming. In addition to the obligate river dolphins, three other small cetaceans also have representatives living in rivers. A subspecies of the finless porpoise (Neophocaena phocaenoides) is endemic to the Yangtze River. Populations of the Irrawaddy dolphin (Orcaella brevirostris) inhabit rivers in Southeast Asia. A subspecies of tucuxi (Sotalia fluviatilis) inhabits rivers in South America, more than 1000 km from the coast. Ranging patterns are highly variable across the dolphins and porpoises. Some of the pelagic animals, such as spinner dolphins, may range over thousands of square kilometers of open ocean, but where they occur near oceanic islands such as Hawaii, they may be locally resident for decades. Bottlenose dolphins may travel hundreds of kilometers from one seamount or island to the next off Bermuda, or during seasonal migrations along the Atlantic seaboard. However, locally resident populations have been reported from many bays, sounds, and estuaries around the world; one such population has been observed over the past 37 years and across five generations in Sarasota Bay, Florida. Similarly, different groups of killer whales near Vancouver Island, Canada, exhibit patterns of long-term residency versus transience; these ranging patterns are



correlated with feeding ecology (see below). Channels and associated habitat limit ranging patterns for river dolphins. One of the smallest documented ranges for a dolphin species was reported for Franciscana dolphins off Argentina, where tagged individuals remained within an area of about 50 km average diameter.

Feeding Ecology In general, dolphins and porpoises tend to eat fish and/or invertebrates (especially squid), that they can swallow whole or after breaking it into smaller pieces. Their jaws and dentition are designed to capture but not chew prey. The size and number of their teeth relate to the size and kinds of prey eaten. Longbeaked dolphins with numerous tiny teeth (e.g., Stenella spp., Delphinus spp., Pontoporia blainvillei) tend to eat small fish and squid, while dolphins with shorter beaks and fewer and larger teeth will take larger prey. The long beaks of river dolphins facilitate obtaining prey in an obstacle-filled environment. At the other extreme, killer whales, with large, wellanchored teeth, can remove pieces from prey much larger than themselves. Further specialization occurs in Risso’s dolphins, where teeth are absent in the upper jaw, and only 10 teeth occur in the lower jaw, for grasping soft-bodied squid. As a departure from all other dolphins and porpoises, the boto (Inia spp.) is the only modern dolphin with differentiated dentition. In the front half of the jaws, the teeth are conical, while further back the teeth have a flange on the inside of the crown, reminiscent of molars, presumably for crushing items from their diverse diet of fish, crabs, and turtles. Different species typically specialize in particular parts of the water column. For example, pantropical spotted dolphins (Stenella attenuata) feed on epipelagic prey, right whale dolphins (Lissodelphis spp.), spinner dolphins (Stenellla longirostris), and Dall’s porpoises (Phocoenoides dalli) feed on mesopelagic prey, while coastal dolphins and porpoises often take advantage of demersal prey. Within species, different age, sex, or social groups may specialize on different prey. For example, lactating dolphins may feed on different species and size classes of prey than dolphins in other physiological states, presumably reflecting different energy and hydration requirements. Cultural differences may also occur. Near Vancouver Island, one group of killer whales specializes on fish such as salmon, while another specializes on marine mammals as prey. Prey distribution influences feeding strategies and behaviors. When prey is relatively evenly distributed, or associated predictably with accessible benthic or

157

shoreline features, then individuals tend to feed alone, as is the case for most porpoises and river dolphins and many coastal dolphins. As prey become more patchy and less predictable, as with schooling fish or squid, then cooperative foraging can be advantageous to dolphins. Schools of dolphins (e.g., Lagenorhynchus spp., Delphinus spp., and Stenella spp.) can spread over relatively large areas to locate rich food patches, and then converge on the prey and work together, circling and concentrating prey schools in order to facilitate prey capture by each individual. Rough-toothed dolphins (Steno bredanensis) have been observed to work in pairs to break large fish into smaller pieces. Long-term social relationships likely facilitate cooperative feeding patterns. Killer whales, with their multigenerational matrilineal social groups, exhibit an extreme form of this behavior, when they work together much as a wolf pack to subdue large prey items such as baleen whales. In a variant on this theme, bottlenose dolphins and Irrawaddy dolphins (Orcaella brevirostris) have learned to work cooperatively with some net fishermen, driving fish toward the fishermen and presumably increasing their own feeding success through the ensuing confusion of the fish or limitations to their movements by net barriers.

Sensory Systems and Communication The aquatic medium limits the utility of some mammalian sensory systems, and enhances others. Olfaction has been reduced in dolphins and porpoises, but some level of chemoreception may be possible. Dolphins and porpoises tend to be very tactile animals, especially in terms of physical contact with conspecifics. Water clarity hampers the use of vision by some dolphins and porpoises in their natural environments, but most that have been tested are believed to have reasonably good vision. However, the eyes of the susu (Platanista gangetica spp.) lack a crystalline lens, making them essentially blind. The density of the aquatic medium provides optimal conditions for sound transmission, and both dolphins and porpoises are highly adapted to take advantage of this feature. Both groups produce echolocation clicks over a broad range of frequencies. This sophisticated system of sound production and echo reception presumably facilitates orientation and navigation, and prey finding in murky or dark waters. Echolocation is also believed to be used for group cohesion by some dolphins. Most dolphins produce pure tone whistles but porpoises do not. A wide variety of whistles are produced in many species. Some dolphins, such as


158


bottlenose (Tursiops spp.), produce individually specific whistles referred to as signature whistles. Playback experiments indicate that these whistles function as long-term individual identifiers recognized by kin and close associates. The distinctive information content appears to reside in the whistle structure itself, rather than being associated with any ‘voice’ features of the individual. Such identifiers may be important for maintaining contact within murky estuarine waters. Dolphins also produce burst-pulse sounds (squawks), primarily in social contexts. It is likely that dolphin social sounds are limited to signaling, emotive, and recognition functions, rather than forming a true language. Killer whales exhibit pod-specific dialects of burst-pulse calls.

Figure 8 Hawaiian spinner dolphin performing the characteristic spinning leap from which its name is derived. Photo by Randall S. Wells.

Behavior and Social Organization Dolphins and porpoises spend more than 95% of their time out of sight below the surface of the water. The behaviors and leaps exhibited at or above the surface by some species represent a small fraction of their full repertoire, which varies by species. With the exception of Dall’s porpoises, which swim rapidly, creating a distinctive ‘rooster tail’, and frequently bowride vessels, most porpoises rarely leap or approach boats. Many dolphins will ride in the pressure waves created by vessels, and some exhibit speciesspecific characteristic leaps, such as the unique multirevolution spins of spinner dolphins (Figure 8), or the high, arcing leaps of pantropical spotted dolphins (Figure 9). Murky waters tend to limit observations of the behavior of river dolphins, but susus (Platanista spp.) have been reported to swim on their side, with their right flipper near the river bottom, echolocating. Feeding behaviors can be dramatic. Killer whales and bottlenose dolphins sometimes briefly beach themselves in pursuit of prey on or near the shore. They will also use their flukes to stun prey, a behavior referred to as ‘fish-whacking’ for bottlenose dolphins. The cooperative efforts of dolphin groups to corral fish schools and drive them to the surface can lead to a feeding frenzy involving large numbers of sea lions and diving birds as well (Figure 10). Some dolphins in Australia carry sponges on their rostra, and are believed to use these as tools for obtaining prey. Some of these feeding behaviors are believed to provide evidence for cultural transmission of knowledge as they are passed between generations and spread across social units. Dolphins and porpoises vary greatly in their degree of sociality. Porpoises and river dolphins tend to live alone or in small groups, with the only persistent groupings being mothers with their most recent

Figure 9 Characteristic high leap by a pantropical spotted dolphin. Photo by Randall S. Wells.

Figure 10 ‘Feeding frenzy’ involving common dolphins, sea lions, pelicans, and gulls, in the Gulf of California. Photo by Randall S. Wells.



calves. All dolphins are social to some degree, ranging from a few individuals to thousands in a school. Oceanic dolphins such as spinner and spotted dolphins tend to form large (sometimes hundreds to thousands) and fluid schools, although some associations within these schools may be long term or recurrent. Some spinner dolphins inhabiting waters near isolated atolls maintain very stable groupings over time. Even greater levels of stability are exhibited by the permanent matrilineal pods of killer whales, which represent several generations of related individuals and remain unchanged except by birth, death, or rare separations. Other large dolphins such as pilot whales are also believed to maintain similar stability, based on results of genetic studies. At an intermediate level, bottlenose dolphins such as those in Sarasota Bay, Florida, may form long-term, multigenerational resident communities of about 150 individuals, where group composition is fluid (fission-fusion), but the site fidelity of the animals guarantees repeated contact as they move through the community range. Longer-term associations within the community include mother–calf bonds lasting 3–6 years, and strong male pair bonds that can persist for decades. In Shark Bay, Western Australia, strong bonds among adult male dolphins translate into complex patterns of alliances of males that work together against other alliances to obtain access to females. At the next level, inter-specific associations can be found among the dolphins, with some of the most common being between spinner and pantropical spotted dolphins, bottlenose dolphins and pilot whales, and Pacific white-sided dolphins, Risso’s dolphins, and northern right whale dolphins. The reasons for these associations are not entirely clear, but they may relate to foraging or protection from predators.

Conservation Status and Concerns Dolphin and porpoise populations around the world typically number in the thousands to millions depending on species and stocks. However, of the 44 species listed in Table 1, as of 2003 six were considered by the IUCN to be critically endangered (n ¼ 2), endangered (n ¼ 2), or vulnerable (n ¼ 2). Many other species lack sufficient information to make any assessment of their status. In each case where concern has been expressed, human activities have been identified as the primary cause for reductions in abundance. It is likely that the critically endangered baiji, or Yangtze River dolphin, has been recently driven to extinction, based on a 2006 survey of the remaining habitat of the species during which no baiji were

159

found (Figure 6). This species was described by Western scientists in 1918. At that time it was still common from Three Gorges to Shanghai. Beginning in 1958, it was hunted intensively for meat, oil, and leather. The decline was exacerbated by accidental losses in fishing activities, including nets, electrofishing and the use of explosives, and entanglement in illegal rolling hooks. Others died from collisions with vessels in the heavily trafficked river. Pollution from the activities of the billion people served by the river, overfishing, underwater blasting for construction, dredging, and damming of tributaries all likely contributed to the decline of this species. The other river dolphins are considered to be endangered (Platanista spp.) or vulnerable (Inia spp.), as a result of their limited habitat, and their close proximity to human activities that directly or indirectly impact the animals, such as habitat fragmentation from damming and fishing. The finless porpoises (Neophocaena phocaenoides) that inhabit the Yangtze River have shown recent significant declines as well. The vaquita, found only in the Gulf of California, is the other of the two critically endangered dolphins or porpoises. Although this porpoise species was only first described in 1958, numbers of remaining vaquita or have been reduced to below 500 individuals due to fishing activities, especially from netting for an endangered sea bass, the totoaba (Figure 7). Found only in the waters of New Zealand, Hector’s dolphins (Cephalorhynchus hectori) have been dramatically reduced in abundance in recent years from fishing activities, and are now considered to be endangered. In addition to the vaquita, there are concerns for all of the porpoises except the spectacled porpoise (Phocoena dioptrica). The harbor porpoise (Phocoena phocoena) is considered vulnerable due to incidental mortalities in fishing nets. Directed fisheries for some dolphins and porpoises occur in many places around the world. For example, striped dolphins (Stenella coeruleoalba) are taken by drive fisheries in Japan, pilot whales (Globicephala melas) are obtained similarly in the Faroe Islands, and other species are obtained by netting or harpoons elsewhere, usually for human consumption, or for bait for other fisheries. One of the largest directed fisheries is prosecuted by the Japanese for Dall’s porpoises (Phoceonoides dalli). The rationale offered for this ongoing fishery is that it provides meat for Japanese markets, partially offsetting the reduction in availability of whale meat from the international whaling moratorium. As indicated above, incidental mortality in fisheries is one of the most serious threats faced by


160


dolphins and porpoises around the world. In the most dramatic example, millions of dolphins of the several species, but especially Stenella attenuata and Stenella longirostris, have been killed in the tuna seine net purse seine fishery in the ETP since the late 1950s. Modifications to fishing practices and gear since the 1990s have reduced annual mortalities to a few thousand individuals, but stocks are not recovering at expected rates, suggesting impacts other than direct mortality in the nets. Incidental mortality is experienced in a wide range of commercial and recreational fisheries around the world. It is managed in some countries, but in many there are no effective controls over takes of marine mammals. Commercial live captures of dolphins for public display, research facilities, and military uses occur in several countries, which supply the rest of the world. Removals of these animals from the wild render them ecologically dead to their native stocks. Typically, these fisheries are prosecuted in the absence of valid scientific assessments of the stocks prior to removals, placing the future for the remaining animals in jeopardy. Institutions in some countries such as the United States have established successful captive breeding programs to avoid these problems, but current breeding programs elsewhere are unable to keep pace with the demands for dolphins for burgeoning programs of interactions with humans around the world, creating pressure for continuing live capture operations. Environmental contaminants pose a highly insidious and likely underappreciated threat to dolphins and porpoises around the world. Very high concentrations of contaminants that cause health and reproductive problems in other mammals have been measured in dolphin and porpoise tissues. Although direct cetacean mortalities from contaminants have not been frequently identified, the weight of evidence is mounting relating concentrations of toxic chemicals to health problems and reproductive failure. Because of the odontocetes’ acoustic sensitivity, underwater noise is also of increasing concern for marine mammals around the world. Noise from vessel traffic, marine construction and demolition, petroleum exploration and production, and military sonars has been implicated in behavioral changes of dolphins and porpoises, and in some cases is believed to have killed cetaceans.

Glossary Anthropogenic Of human origin, as in man-made threats to dolphins or porpoises, such as pollution,

fishing gear, or noise from industrial or military activities. Anti-tropical Distributed to the north and to the south of the tropics. Biotoxin Naturally created toxin produced by organisms such as algae. By-catch Organisms caught unintentionally in the course of catching target species. Cetacean Whale, dolphin, or porpoise, of the mammalian order Cetacea. Demersal Near the seafloor, as in prey fish that tend to not move through the water column or to the surface of the water. Drive Fishery Fishery directed at cetaceans in which noise from an array of small boats is used to move the animals onto a beach. Echolocation (sonar) Sophisticated acoustic orientation system of odontocetes in which clicks produced in air passages are focused and projected forward, and received echoes from the clicks provide information about the cetacean’s environment. Epipelagic Near the water’s surface, as in prey fish that do not inhabit the lower portions of the water column or the seafloor. Mesopelagic Within the water column, as in prey fish that are not commonly found near the surface or the seafloor. Odontocete A toothed cetacean, of the suborder Odontoceti. Paedomorphosis The retention of juvenile characteristics in adults. Pantropical Found in tropical waters throughout the world. Peduncle The tail stock, or connection between the body and the flukes of a cetacean. Pelagic Open ocean waters, typically deep and far from shore. Rostrum The beak or tooth-filled projection of the mouth of a cetacean. Sexual dimorphism Differences in the body form or size based on gender. Strandings When live or dead cetaceans become beached.

See also Marine Mammal Diving Physiology. Marine Mammal Evolution and Taxonomy. Marine Mammal Migrations and Movement Patterns. Marine Mammal Social Organization and Communication. Marine Mammal Trophic Levels and Interactions. Marine Mammals and Ocean



Noise. Marine Mammals, History of Exploitation. Pollution: Effects on Marine Communities. Sonar Systems.

Further Reading Berta A and Sumich JL (2003) Marine Mammals: Evolutionary Biology. San Diego, CA: Academic Press. Leatherwood S and Reeves RR (eds.) (1990) The Bottlenose Dolphin. San Diego, CA: Academic Press. Mann J, Connor RC, Tyack PL, and Whitehead H (eds.) (2000) Cetacean Societies: Field Studies of Dolphins and Whales. Chicago, IL: University of Chicago Press. Norris KS, Wu¨rsig B, Wells RS, and Wu¨rsig M (1994) The Hawaiian Spinner Dolphin. Los Angeles, CA: University of California Press. Read AJ, Wiepkema PR, and Nachtigall PE (eds.) (1997) The Biology of the Harbour Porpoise. Woerden: De Spil Publishers. Reeves RR, Smith BD, Crespo EA, and Notarbartolo di Sciara G (eds.) (2003) Dolphins, Whales, and Porpoises: 2002–2010 Conservation Action Plan for the World’s Cetaceans. Gland: IUCN/SSC Cetacean Specialist Group. Reeves RR, Smith BD, and Kasuya T (eds.) (2000) Biology and Conservation of Freshwater Cetaceans in Asia. Gland: IUCN.

161

Reynolds JE, III, Perrin WF, Reeves RR, Montgomery S, and Ragen TJ (eds.) (2005) Marine Mammal Research: Conservation Beyond Crisis. Baltimore, MD: The Johns Hopkins University Press. Reynolds JE, III and Rommel SA (eds.) (1999) Biology of Marine Mammals. Washington, DC: Smithsonian Institution Press. Reynolds JE, III, Wells RS, and Eide SD (2000) The Bottlenose Dolphin: Biology and Conservation. Gainesville, FL: University Press of Florida. Rice DW (1998) Special Publication No. 4: Marine Mammals of the World: Systematics and Distribution. Lawrence, KS: The Society for Marine Mammalogy. Ridgway SH and Harrison R (eds.) (1989) Handbook of Marine Mammals, Vol. 4: River Dolphins and the Larger Toothed Whales. San Diego, CA: Academic Press. Ridgway SH and Harrison R (eds.) (1994) Handbook of Marine Mammals, Vol. 5: The First Book of Dolphins and Porpoises. San Diego, CA: Academic Press. Ridgway SH and Harrison R (eds.) (1999) Handbook of Marine Mammals, Vol. 6: The Second Book of Dolphins and Porpoises. San Diego, CA: Academic Press. Twiss JR, Jr. and Reeves RR (eds.) (1999) Conservation and Management of Marine Mammals. Washington, DC: Smithsonian Institution Press.


DOUBLE-DIFFUSIVE CONVECTION R. W. Schmitt, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction The density of seawater is determined by both its temperature and its salt content or salinity. Whereas added heat makes water lighter, added salt makes it denser, so both must be considered when evaluating the gravitational stability of the water column. That is, a given column of water will ‘convect’ or overturn if dense waters overlie lighter waters. In many parts of the world ocean, the distributions of temperature and salinity are opposed in their effects on density. This arises because of the tendency of warm water to easily evaporate in low latitudes, the predominance of rainfall in cold, high-latitude regions, and the deep circulation patterns that bring the cold waters to lower latitudes. The opposing effects of temperature and salinity on density, and the fact that the molecular conductivity of heat is about 100 times as large as the diffusivity of salt in water, makes possible a variety of novel convective motions that have come to be known as double-diffusive convection. In the following, the oceanic double-diffusive mixing phenomena such as ‘salt fingers’, ‘diffusive convection’, and ‘intrusions’ are discussed in turn. Observational evidence suggests their importance in all the oceans, and models indicate a substantial impact on water mass structure and the thermohaline circulation.

Salt Fingers In much of the subtropical ocean, warm, salty water near the surface overlies cooler, fresher water from higher latitudes. If the temperature contrast could be removed there would be a large-scale overturning of the water column, releasing the very substantial energy available in the salt distribution. However, this does not happen except on a small scale, where the greater diffusivity of heat can establish thermal equilibrium in adjacent water parcels that still have strong salt contrasts. A bit of warm, salty water displaced into the cold freshwater beneath loses heat, but not much salt to the surrounding water, leaving a cool, salty water parcel that continues to sink. Similarly, a cold fresh parcel displaced upward gains

162

heat but not salt, becoming warm and fresh and therefore buoyant. This ‘salt finger’ instability, discovered by M. Stern in 1960, appears as a closepacked array of up and down flowing convection cells which exchange heat laterally but diffuse little salt. The result is an advective transport of salt and, to a lesser extent, heat in the vertical. Typical cell widths in the ocean are 2–3 cm, the scale for effective heat conduction. The salt finger instability is ‘direct’, in the sense that initial displacements are accelerated, and can be modeled accurately with an exponential growth rate. When most intense, the fingers tend to exist on high-gradient interfaces separating wellmixed layers in the adjacent fluid. The significant role of salt fingers in oceanic mixing is now becoming apparent, as there are clear indications that it is the dominant mixing process in certain regions and a contributing process within the main thermocline of the subtropical gyres. As could be expected, the propensity toward salt fingering is a strong function of the intensity of the vertical salinity gradient. The instability can grow at extremely weak values of the salinity gradient, because the diffusivity of salt is 2 orders of magnitude less than the thermal conductivity. When expressed in terms of the effects on density, all that is required is a top-heavy density gradient due to salt that is only about one-hundredth of the gradient due to temperature. That is, the density ratio, Rr, must be less than the diffusivity ratio: 1oRr aTZ =bSZ okT =kS E100

½1

where a, b are the thermal expansion and haline contraction coefficients, TZ, SZ are the vertical gradients of temperature and salinity, and kS, kT are the molecular diffusivity for salt and the thermal conductivity. This criterion is met over vast regions of the tropical and subtropical thermocline, since the ratio of diffusivities is about 100. However, while the required salt gradient is very small, the growth rate of salt fingers does not become ‘large’ until Rr approaches 1 (at which the vertical density gradient vanishes). Indeed, we find that the primary fine-scale evidence for salt fingers, the ‘thermohaline staircase’, occurs only when the density ratio becomes low. Fingers transport more salt than heat in the vertical and have a net counter-gradient buoyancy flux. Since the growth rate and fluxes increase with the strength of the stratification, high-gradient regions will harbor greater fluxes than adjacent weak-gradient intervals.


DOUBLE-DIFFUSIVE CONVECTION

This leads to a buoyancy flux convergence that can cause the weaker-gradient region to overturn and mix. The resulting structure has thin interfaces separating thicker, well-mixed layers. The layers are continuously mixed by the downward salt flux, and the convective turbulence of the layers serves to keep the interface thin and limits the length of the fingers. Observations of the ‘thermohaline staircase’ have been reported from several sites with strong salinity gradients. A necessary condition for an organized salt finger staircase seems to be that the density ratio is less than 1.7 (Figure 1). Such conditions are found occasionally near the surface, where evaporation produces the unstable salinity gradient, but more often at depth where the presence of isopycnal

gradients of temperature and salinity can lead to a minimum in Rr, provided there is a component of differential advection (shear) acting on the isopycnal gradients of T and S. Examples of staircases are found beneath the Mediterranean water in the eastern Atlantic, within the Mediterranean and Tyrrhenian Seas, and beneath the subtropical underwater (salinity maximum) of the western tropical Atlantic. Detailed examinations of one particular staircase system in the western tropical Atlantic were made in 1985 (Caribbean Sheets and Layers Transects – C-SALT) and in 2001, when a ‘Salt Finger Tracer Release Experiment’ (SFTRE) was performed. Over a large area in the western tropical North Atlantic (B1 million km2), a sequence of B10–15 mixed

Mediterranean outflow, R 1.13

Tyrrhenian Sea, R 1.15

3 4

Pressure (db)

800

5

900

1

2

2 3 4

Station 19

5

6

7

8

1100

8

1100

66

9

10

1400

10

1500

1500

T

1540 A II 76 SCIMP 2 17 VII 73

1300

1300

70

1530

1550

9

68

S

1520

1200

1200

1400

800

1000

64

1510

700

900

6

7

1000

62

Depth (db)

1

700

Temperature (C)

Salinity (‰) 3850 3870

Temperature (C) 13 132013401360

163

1560 3540

3545

3550

3555

Salinity (‰) Subtropical underwater R 1.6 Temp (C) 5 Salt (‰) 33 100

10 34

15 35

N.A. central water R 1.9

20 36

25 37

Salt (‰) Temp (C)

3632 3538 3544 3550 3556 1000 1025 1050 1075 1100

760 T

Pressure (db)

200

300

400

S

S Pressure (db) CTD 26

T

820

T

880 T 940

S T

500

1000

Figure 1 The occurrence of thermohaline staircases as a function of density ratio. The strong staircases have Rro1.7. Irregular steppiness characterizes the central waters of the subtropical gyres, where RrB2.


164


layers, 5–40-m thick, can be observed. Data from the 1960s to the 2000s indicate that the layers are a permanent feature of the region, despite layer splitting and merging, and a moderately strong eddy field. One of the most remarkable characteristics of this staircase is the observed change in layer properties across the region. Layers get colder, fresher, and lighter from north to south (Figure 2), the inferred flow direction for the upper layers, which appear to be losing salt to the layers below. These unique water mass transformations in temperature, salinity, and density provide strong evidence for salt fingers. That is, such changes can only be due to a flux convergence by salt fingers, which transport more salt than heat; turbulence transports the two components equally and isopycnal mixing, by definition, transports them in density-compensating amounts. Towed microstructure measurements taken in the staircase revealed limited-amplitude, narrow-band temperature structure within the interfaces. The dominant horizontal wavelength was B5 cm, in excellent agreement with the theoretical finger scale. The shape of towed microstructure spectra for this and many other observations is also found to be distinctly different from that of turbulence, leading to useful discrimination tests for towed data.

Vertically profiling instruments which measure the dissipation rates of both thermal variance and turbulent kinetic energy reveal a strong correlation of thermal variance with the interfacial gradients that provide the strongest finger growth rate (Figure 3). Such data also allow the discrimination of salt fingers from turbulence by their relative efficiencies in converting energy sources into changes in the stratification. That is, salt fingers are rather efficient in converting energy from the salt field to the thermal field (B70%), with the result that there is relatively little viscous dissipation for the amount of mixing achieved. In contrast, turbulence is rather inefficient, with only B20% of dissipated kinetic energy converted into an increase in potential energy. Also, salt fingers lead to a net decrease in potential energy, exactly the opposite to turbulence (this is what allows it to maintain staircase-type profiles, whereas turbulence should ultimately smooth the overall profiles). A good way to appreciate this difference in mixing mechanisms is to compare the formulas for estimating the vertical diffusivities from microstructure measurements of the dissipation rates of turbulent kinetic energy (e) and thermal variance (w). These formulas are contrasted below:

•

for turbulence (with flux Richardson number, Rf ¼ 0.1770.03, after Osborn (1980) and Osborn and Cox (1972)):

1.8

14

1.6

12

10

1.2

27.2 26.8

1 27.4

0.8

27

0.6

9

0.4

8 7 34.8

Ky ¼ KS ¼

1.4

13

11

Ky ¼ KS ¼ Kr ¼

log10 scale

Potential temperature (C)

Volumetric TS diagram 15

0.2 35

35.2

35.4 35.6 Salinity

35.8

36

0

Figure 2 Potential temperature–salinity values from a depth cycling CTD (an acronym for ‘conductivity, temperature, and depth’) on a mooring in the center of the tropical Atlantic thermohaline staircase during SFTRE. The color intensity represents the number of observations of any one T–S value; thus the mixed layers appear as distinct high-density lines in this 4.5-month time series. The evolution of layer properties across the region is such that layers become warmer, saltier, and denser from southeast to northwest, and the advection of the layers past the mooring provides the range of T–S values observed. The layer properties cross isopycnals (the 26.8, 27.0, 27.2, and 27.4 potential density surfaces are shown) with an apparent heat/salt density flux convergence ratio near 0.85.

•

Rf e e e ¼ Gt 2 E0:2 2 1 Rf N2 N N

wy 2yZ

½2

for salt fingers (with Rr ¼ 1.6, and flux ratio g ¼ 0.7, after St. Laurent and Schmitt (1999)): Rr 1 e e E2 2 1 g N2 N Rr w y KS ¼ E2:3Ky g 2yZ w g g Rr 1 e Ky ¼ y ¼ KS ¼ 2yZ Rr Rr 1 g N 2 e e ¼ Gf 2 E0:8 2 N N

KS ¼

½3

Note that the ‘mixing efficiencies’ Gt, Gf are distinctly different for turbulence and salt fingers, with the fingers being more efficient and dissipating less energy for a given amount of mixing. A broad-scale microstructure survey capable of addressing these issues was done with the North Atlantic Tracer Release Experiment (NATRE;



34.8

200

200

250

250

300

300 T

S

350

350

400

400

450

450

500

8

10 12 14 Temperature (C)

16

109

107 105 106 108 Dissipation of thermal variance (K2 s1)

Depth (m)

Depth (m)

Salinity 35.2 35.6

165

500 0.001 0.002 0.003 0 Salt finger growth rate (s1)

Figure 3 Profiles of potential temperature and salinity (left panel), the dissipation rate of thermal variance (middle panel), and the theoretical salt finger growth rate (right panel) for a high resolution profiler cast during SFTRE. The tracer was injected into the layer with potential temperature near 10 1C (B350-m depth). The salt finger growth rate calculated from the fine-scale temperature and salinity gradients is a good predictor of the microscale dissipation rate of thermal variance.

see Tracer Release Experiments). This region of the eastern North Atlantic thermocline is susceptible to salt fingers, and optical microstrucuture imagery revealed that they were the most frequently observed microstructure in the thermocline. In addition, it was obvious in the sensor data that there were many occurrences of the ‘high-chi, low-epsilon’ signature of salt fingers, that contrasted with the high-epsilon signatures of turbulence. A parametric sorting of the mixing events allows classification of the stronger microstructure patches by the value of the local Richardson number and density ratio. Statistically significant variations in the value of the ‘mixing efficiency’ were observed in this parameter space (Figure 4). When translated into a flux ratio for salt fingers, the oceanic microstructure data are in excellent agreement with laboratory experiments on salt fingers. This parametric approach to the microstructure allows a classification of the mixing events as either turbulent (with low efficiency) or salt fingering (with high efficiency). With each occurring at different

frequencies in the water column, this translates into differences for the net vertical eddy diffusivities for heat and salt. At the depth range of the tracer injection in NATRE, the diffusivity estimated taking salt fingers into account agrees well with the value derived from tracer dispersion. Analysis using the conventional turbulence formula yields a diffusivity that is 50% low and a diapycnal velocity of the wrong sign. The magnitude of the fluxes in NATRE can be contrasted with those in the C-SALT/SFTRE staircase. From the substantial rate of dissipation of thermal variance, we can estimate an eddy diffusivity for salinity of 0.9 10 4 m2 s 1 within the staircase. This is in excellent agreement with the observed dispersion of tracer within the staircase, showing that diffusivities are elevated by an order of magnitude by the formation of steps. Since the staircase occupies about one-fourth of the area of the Atlantic between 101 and 151 N, the vertical salt flux in this high gradient area is predicted to be 3–4 times as large as the flux in the remaining area of this latitude band. This is because the rest of the area is expected to have a


166

DOUBLE-DIFFUSIVE CONVECTION 0

0.1

0.2

0.3

0.4

0.5

0.6

Doubly stable regime

0.7

0.8

1

R

R 100 30 10 5 2.5 1.5 2

0.9

Finger-favorable regime

1.1

1.01

1.01

1.1

1.5

2.5

5

10

30

100 100

log10 Ri

2 1

0

Ri

10 5

1

0.25 0.1

1

2

0.01 2

1

0 log10 (IR I1)

1

2

2

1

0 log10 (R 1)

1

2

Figure 4 The ‘mixing efficiency’, G, as a function of density ratio and Richardson number, for the microstructure sampled in the main thermocline of the eastern North Atlantic Ocean (right panel). The left panel shows data from non-double-diffusive regions, where results are consistent with turbulence. The low-Rr, high-Ri regime of the right panel provides clear indications of high-efficiency salt fingers playing a significant role in the mixing of the North Atlantic thermocline.

diffusivity 10 times smaller (like NATRE) as well as a weaker salinity gradient. Thus, the staircase areas appear to be very significant sites of enhanced diapycnal exchange and water mass transformation.

Diffusive Convection The ‘diffusive’ form of double-diffusive convection is realized when the stratification is the opposite of the salt finger situation. That is, cold fresh water overlies warm salty water, with the salt providing the overall stabilization of the water column. However, there is energy to be released in the ‘warm on the bottom’ temperature distribution, and the different rates of heat and salt diffusion allow convection to occur. The essential physics is distinct from salt fingers, as the faster diffusion of heat is releasing energy in its own distribution rather than that of the slowerdiffusing salt. Again considering movement of small parcels of water, we see that the elevation of warm salty water into the cold fresh one will cause it to become cold salty water, and thus heavier than when it started upward. Instead of accelerating upward as in a salt finger, it is actually driven back down with greater force than it took to initially displace it. This is termed an ‘overstability’ and leads to a growing

oscillation. However, the oscillatory behavior is hard to observe except in careful laboratory experiments, as it quickly reaches an amplitude where transition to a layered series of convective cells is realized. This is another form of thermohaline staircase, with temperature and salinity both increasing with depth. A laboratory experiment that involves heating a stable salt gradient from below easily develops a thermohaline staircase in the diffusive sense. The mixed layers are maintained by convective motions driven by the heat flux from below; the thin, stable, gradient regions are sharp interfaces that conduct heat vertically, but transport little salt. In the ocean such ‘diffusive’ staircases are mostly found in highlatitude oceans, where surface cooling and freshening can set up the necessary gradients. Often, diffusive staircases are found under sea ice (Figure 5). It seems that the ice helps to isolate the water column from wind forcing, leading to exceptionally weak internal waves, so that the relatively slow diffusive process can dominate the vertical mixing. The fluxes for the diffusive staircase are generally less than fluxes for salt fingers. This is because fingers advectively carry heat and salt vertically across the interfaces, whereas a diffusive interface must rely largely on vertical conduction across the horizontal interface. The surface area for heat diffusion is much



Temperature (°C) 1.0 0

150

300

310

Depth (m)

250 320 (b) Section of profile recorded at high gain

Depth (m)

(a) Typical temperature profile section 200

300

330 350

340 0.01 °C 0.1 C Figure 5 Temperature staircase in the Arctic halocline beneath the ice. Five- to ten-meter mixed layers are seen separated by thin ‘diffusive’ interfaces across which heat conduction provides a buoyancy flux to stir the adjacent layers. Such staircase layers are widespread beneath the ice and the upward heat flux they supply may contribute to its melting.

greater in the convoluted structure of a salt finger interface. In many ways, the fully developed diffusive interface is simply a modified form of Rayleigh– Bernard convection, with fluid boundary conditions and the diffusion of salt acting as a weak drag on the intensity. The relative effectiveness of heat and salt diffusion across the interface sets the salt to heat flux ratio. Except for density ratios very close to 1, the flux ratio is rather low. This can be understood by considering an interface made sharp by convection. The heat and salt would diffuse into boundary layers on either side, with different thicknesses depending on the square root of their diffusivities. After a certain time, the thermal boundary layer would be thick enough to be unstable and convection would occur, carrying the heat and salt anomalies away from the interface. The relative amounts of salt and heat transported should depend on the square root of the diffusivity ratio (B0.1), a number in reasonable agreement with laboratory measurements, so long as the density ratio is not too close to 1. At density ratios closer to 1, the interface is increasingly disrupted by turbulent plumes from the mixed layers, and a more direct transport of both heat and salt occurs, resulting in a higher salt-to-heat buoyancy flux ratio. Of course, this ratio is limited by the energetics to be less than 1.

167

Strong but localized diffusive staircases are found at the hot, salty brines in topographic deep areas found at oceanic spreading centers. There the separation of heat and salt could be contributing to pooling of the brines in topographic depressions and possibly ore formation as well. The diffusive process also plays a role in intrusions and on the upper side of warm, salty water masses such as the Mediterranean water in the Atlantic. It may be a factor in the evolution of fresh mixed layers laid down by rain, river inputs, or ice melt. These can create ‘barrier layers’ whereby the strong salt stratification prevents mixing with underlying water. If the fresh surface layers cool, the conditions for diffusive convection arise at the base of the mixed layer. Since diffusive convection transports heat upward, but not much salt, there is a tendency to maintain the barrier layer stratification. Thus, such barrier layers should persist longer than they would without double diffusion. Barrier layers appear to be important in modifying air–sea interactions in both the Tropics and highlatitude areas of deep convection such as the Labrador Sea. However, the most extensive regions of diffusive convection are found beneath the surface layers in the polar and subpolar oceans. Steps under the Arctic ice were first reported in the 1970s, and have been observed to cover much of the Arctic in recent data. A ‘diffusive’ thermohaline staircase between about 200- and 400-m depth appears to be a ubiquitous feature under most of the Arctic ice field away from the boundaries. It supports a heat flux from the intruding warm, salty Atlantic water to the cooler, fresher Arctic surface waters above. The extensiveness of the staircase may be due to the especially weak internal wave field under the ice. This is likely due to the rigid ice lid, but also possibly due to an enhanced wave decay within the convectively mixed staircase itself. Areas near topography with stronger internal waves and more frequent turbulent mixing events are less likely to harbor a staircase. The downgradient buoyancy flux from the turbulence (an upgradient buoyancy flux is necessary to maintain a staircase) and the destruction of the small-scale property gradients by isotropic turbulence are competing factors to the double diffusion. The waters around Antarctica also display prominent diffusive staircases. The layers in the Weddell Sea are much thicker (10–100 m) than those found in the Arctic and may support an upward heat flux of 15 W m 2 in open waters, if the diffusive interfaces are thin enough, an issue not easily resolved with ordinary instruments. This flux is sufficient to be important in upper ocean heat budgets and may help to maintain ice-free conditions in the summer. Lower


168


fluxes are estimated for the Arctic steps. However, recent observations of a much stronger incursion of Atlantic water into the Arctic are characterized by very extensive double-diffusive intrusions. This suggests that the lateral processes (next section) may be dominant over vertical in the halocline of the Arctic Ocean, and that the significant climatic changes currently underway there may be mediated in part by small-scale double-diffusive processes. The importance of diffusive convection in polar regions lies in its ability to produce a cold, salty, and dense water mass without air–sea interaction. That is, heat can be extracted from a subsurface water mass without much change in salinity. The resulting water may be dense enough to become a bottom water mass. This idea has been applied to the formation of Antarctic Bottom Water and Greenland Sea Bottom Water. The T–S characteristics are in agreement with the model predictions but quantification of the rates of mixing remains uncertain. This is now viewed as a critical issue since the upward heat flux may be contributing to the current rapid decay of Arctic sea ice.

Intrusions In the presence of horizontal variations in temperature and salinity along density surfaces, as is common at oceanic fronts, the small-scale double-diffusive processes can drive horizontal motions on 10–100-m vertical scales. These intrusive instabilities arise because of the buoyancy flux convergences due to the mixing by salt fingers and diffusive interfaces. In the presence of horizontal T and S gradients, these vertical flux convergences generate lateral pressure gradients which drive a slow movement of water across the front. This is often manifested as a complex interleaving of warm/salty and cold/fresh water masses (Figure 6). The relative motion of each water type relative to the other is an effective means for keeping the double diffusion most intense, as it has a tendency to drive the density ratio toward 1. Thus, it is a selfreinforcing (direct) instability or a sort of ‘horizontal salt finger’. The sense of the heat and salt flux convergences is such that a warm salty intrusion is expected to lose more salt than heat due to fingering across its lower boundary, and thus should become lighter and rise across density surfaces. Similarly, a cold fresh intrusion should gain more salt than heat and become denser and sink across density surfaces. Such behavior was predicted theoretically by Stern, confirmed in the laboratory by Turner and observed in numerous observational programs. In situations where the temperature increases with depth, diffusive

convection may dominate the mixing, leading to sinking of a warm salty intrusion. Microstructure observations confirm that enhanced dissipation consistent with double diffusion occurs at the interfacial boundaries of such intrusive fine structure. For more detail on double-diffusive intrusions, see Intrusions.

Global Importance It is fair to ask whether the different heat/salt transport rates achieved in small-scale doublediffusive mixing have any influence on the large-scale circulation. In general, ocean models assume that the small scales are available to consume any necessary variance, and in particular that there is no difference in heat and salt diffusivities. The presence of double diffusion in the ocean means that this assumption is invalid, and that a variety of effects whereby the differential transport rates feedback on the largerscale structure manifest themselves. For the intrusive instabilities described above, one effect is the tendency to destroy small-scale anomalies in temperature and salinity. The role of the relative horizontal motion between the vertically arrayed layers (shear) in forcing the density ratio toward 1 is important here, as this keeps the doublediffusive convection most intense. The strong mixing continues driving the anomalous fluid across density surfaces until it reaches a level with matching properties. Thus, intrusions are a powerful mechanism for removing water-mass anomalies and maintaining the tightness of the mean temperature– salinity relationship. The process can occur anywhere, and there is good evidence that it is a major lateral mixing agent at both polar and equatorial latitudes. Another effect on the T–S relation is due to the strong dependence of the vertical mixing rate on density ratio. The well-documented increase in fingering intensity as the density ratio approaches 1 leads to a number of interesting effects. When vertical variations in density ratio arise this dependence leads to fine-scale flux convergences which act to remove the anomaly in density ratio. Since densitycompensated T–S anomalies are prominent in the mixed layer, such a differential mixing mechanism is needed to explain the tightness and shape of the T–S relation of the subducted waters in the thermocline. Also, for both salt fingers and diffusive convection, there is a forcing of the density ratio away from unity, unless compensated by vertical fluxes or differential lateral advection. In addition, salt fingers are often the dominant mixing process operating on fine-scale intrusions at fronts. Double-diffusive


DOUBLE-DIFFUSIVE CONVECTION 5 26.15

0

169

5 10 Mixing intensity 26.55

Density ref. prs 0.0 26.95 27.35

27.75

Salinity (‰) 33.50

33.90

34.30

34.70

35.10

5.20

6.80

(C) ref. prs 0.0 0.40 0

2.00 S

3.60

T

100

Pressure (db)

200

300

400

500

T

S

600 Figure 6 Intrusive fine structure in the front associated with the North Atlantic Current east of Newfoundland. Warm, salty waters from the south interleave with cold, fresh waters from the north to create strong salinity-compensated temperature inversions with an overall stable density profile (s). An optical microstructure instrument revealed intense double-diffusive mixing at the boundaries of the intruding water masses.

intrusions may be a primary mechanism for accomplishing lateral mixing of water masses at the fine scale, acting efficiently on the horizontal gradients produced by meso-scale stirring by eddies. Since double diffusion acts preferentially on high-gradient regions, it may be responsible for a large fraction of the global dissipation of thermal and haline variance, despite modest eddy diffusivities. This is reinforced by the recent discovery that enhanced open ocean turbulence is found mainly in the weakly stratified abyss, where the contribution to dissipation of scalar variance is necessarily small, even though the eddy diffusivities may be large. In addition to being of regional importance as an enhanced flux site in the thermocline of the tropical

North Atlantic, salt fingers may be important in all of the other oceans and many marginal seas. In the Atlantic at 241 N, fully 95% of the upper kilometer of the ocean is salt finger favorable. Indeed, conditions are favorable for fingering in all of the central waters of the subtropical gyres. Since it is well established that the strength of the thermohaline circulation is very sensitive to the magnitude of the vertical (diapycnal) mixing coefficient, we must be concerned with the large-scale effects of widespread salt fingering. Model studies show that major features of the steady-state solutions are very sensitive to the ratio of the vertical eddy diffusivities for salinity and temperature. A 22% decrease in the strength of the thermohaline circulation was realized


170


in one model when salt fingers were added to the vertical mixing scheme. Studies in global models with realistic topography and forcing found that double diffusion helps to bring deep temperature and salinity fields into closer agreement with observations. Models also suggest that double diffusion lowers the net interior density diffusivity sufficiently to make the thermohaline circulation more susceptible to collapse. Since collapse of the thermohaline circulation has occurred rapidly in the past, has dramatic impacts on climate, and is predicted to be a possible outcome of future greenhouse warming, it behooves us to seek a better understanding of the double-diffusive mixing processes in the ocean.

See also Deep Convection. Intrusions. Open Convection. Tracer Release Experiments.

Ocean

Further Reading Osborn T (1980) Estimates of the local rate of vertical diffusion from dissipation measurements. Journal of Physical Oceanography 10: 83--89. Osborn T and Cox C (1972) Oceanic fine structure. Geophysical Fluid Dynamics 3: 321--345. Schmitt RW (1994) Double diffusion in oceanography. Annual Reviews of Fluid Mechanics 26: 255--285. Schmitt RW, Ledwell JR, Montgomery ET, Polzin KL, and Toole JM (2005) Enhanced diapycnal mixing by salt fingers in the thermocline of the tropical Atlantic. Science 308(5722): 685--688. Stern ME (1975) Ocean Circulation Physics. New York: Academic Press. St. Laurent L and Schmitt RW (1999) The contribution of salt fingers to vertical mixing in the North Atlantic Tracer Release Experiment. Journal of Physical Oceanography 29(7): 1404--1424. Turner JS (1973) Buoyancy Effects in Fluids. Cambridge, UK: Cambridge University Press.


DRIFTERS AND FLOATS P. L. Richardson, Woods Hole Oceanographic Institution, Woods Hole, MA, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction Starting in the 1970s, surface drifters and subsurface neutrally buoyant floats have been developed, improved, and tracked in large numbers in the ocean. For the first time we have now obtained worldwide maps of the surface and subsurface velocity at a few depths. New profiling floats are measuring and reporting in real time the evolving temperature and salinity structure of the upper 2 km of the ocean in ways that were impossible a decade ago. These measurements are documenting variations of the world ocean’s temperature and salinity structure. The new data are revealing insights about ocean circulation and its time variability that were not available without drifters and floats. Surface drifters and subsurface floats measure ocean trajectories that show where water parcels go, how fast they go there, and how vigorously they are mixed by eddies. Ocean trajectories, which are called a Lagrangian description of the flow, are useful both for visualizing ocean motion and for determining its velocity characteristics. The superposition of numerous trajectories reveals that very different kinds of circulation patterns occur in different regions. Time variability is illustrated by the tangle of crossing trajectories. Trajectories often show the complicated relationship between currents and nearby seafloor topography and coastlines. Drifters and floats have been used to follow discrete eddies like Gulf Stream rings and ‘meddies’ (Mediterranean water eddies) continuously for years. When a drifter becomes trapped in the rotating swirl flow around an eddy’s center, the path of the eddy and its swirl velocity can be inferred from the drifter trajectory. Thus the movement of a single drifter represents the huge mass of water being advected by the eddy. Trajectories of drifters launched in clusters have provided important information about dispersion, eddy diffusivity, and stirring in the ocean. When a sufficient number of drifters are in a region, velocity measurements along trajectories can be grouped into variously sized geographical bins and calculations made of velocity statistics such as mean velocity, seasonal variations, and eddy energy. Gridded

values of these statistics can be plotted and contoured to reveal, for example, patterns of ocean circulation and the sources and sinks of eddy energy. Maps of velocity fields can be combined with measurements of hydrography to give the three-dimensional velocity field of the ocean. Oceanographers are using the newly acquired drifter data in these ways and also incorporating them into models of ocean circulation. Care must be used in interpreting drifter measurements because they are often imperfect current followers. For example, surface drifters have a small downwind slip relative to the surrounding water. They also tend to be concentrated by currents into near-surface converge regions. The surface water that converges can descend below the surface, but drifters are constrained to remain at the sea surface in the convergence region. Surface currents sometimes converge drifters into swift ocean jets like the Gulf Stream. This can result in oversampling these features and in gridded mean velocities that are different from averages of moored current meter measurements, which are called Eulerian measurements of the flow. Bin averages of drifter velocities can give misleading results if the drifters are very unevenly distributed in space. For example, drifters launched in a cluster in a region of zero mean velocity tend to diffuse away from the cluster center by eddy motions, implying a divergent flow regime. The dispersal of such a cluster gives important information about how tracers or pollutants might also disperse in the ocean. Drifters tend to diffuse faster toward regions of higher eddy energy, resulting in a mean velocity toward the direction of higher energy. This is because drifters located in a region of high eddy kinetic energy drift faster than those in a region of low kinetic energy. Errors concerning array biases need to be estimated and considered along with gridded maps of velocity.

Surface Drifters Surface drifter measurements of currents have been made for as long as people have been going to sea. The earliest measurements were visual sightings of natural and man-made floating objects within sight of land or from an anchored ship that served as a reference. Starting at least 400 years ago, mariners reported using subsurface drogues of different shapes and sizes tethered to surface floats to measure currents (Figure 1). The drogues were designed to have a large area of drag relative to the surface float so that


171

172

DRIFTERS AND FLOATS

compilations of historical ship drift measurements. Pilot charts used by most mariners today are still based on historical ship drifts. Problems with the ship drift technique are the fairly large random errors of each velocity measurement (B20 cm s 1) and the suspected systematic downwind leeway or slip of a ship through the water due to wind and wave forces. New velocity maps based on satellite-tracked drifters are providing a much more accurate and higherresolution replacement of ship drift maps. Drifting derelict ships gave an early measurement of ocean trajectories during the nineteenth century. Wooden vessels that had been damaged in storms were often abandoned at sea and left to drift for months to years. Repeated sightings of individual vessels reported in the US pilot charts provided trajectories. Other Drifters

Figure 1 Schematic of an early drifter and drogue from the Challenger expedition (1872–76). Adapted from Niiler PP, Davis RE, and White HJ (1987) Water following characteristics of a mixed layer drifter. Deep-Sea Research 24: 1867–1881.

the drifter would be advected primarily with the water at the drogue depth and not be strongly biased by wind, waves, and the vertical shear of nearsurface currents. Over the years many kinds of drogues, tethers, and surface floats have been tried, including drogues in the form of crossed vanes, fishing nets, parachutes, window shades, and cylinders.

Bottles with notes and other floating objects have been a popular form of surface drifter over the years. The vectors between launch and recovery on some distant shore provided some interesting maps but ones that were difficult to interpret. The technique was improved and exploited in the North Atlantic by Prince Albert I of Monaco during the late 1800s. More recently, 61 000 Nike shoes and 29 000 plastic toy animals were accidentally released from damaged containers lost overboard from ships in storms in the North Pacific. The recovery of thousands of these drifters along the west coast of North America has given some interesting results about mean currents and dispersion. Bottom drifters are very slightly negatively buoyant and drift along the seafloor until they come ashore and are recovered. The vectors between launch and recovery show long-term mean currents near the seafloor. Tracking

Ship Drifts

Probably the most successful historical drifter is a ship; the drift of ships underway as they crossed oceans provided millions of ocean current measurements. A ship drift measurement is obtained by subtracting the velocity between two measured position fixes from the estimated dead reckoning velocity of the ship through the water over the same time interval. The difference in velocity is considered to be a measure of the surface current. This technique depends on good navigation, which became common by the end of the nineteenth century. Most of what we have learned about the large-scale patterns of ocean currents until very recently came from

Early measurements of drifters were visual sightings using telescopes, compasses, and sextants to measure bearings and locations. Later during the 1950s, radio direction finding and radar were used to track drifters over longer ranges and times from shore, ship, and airplane. Some drifter trajectories in the 1960s were obtained by Fritz Fuglister and Charlie Parker in the Gulf Stream and its rings using radar. These early experiments did not obtain very many detailed and long trajectories but did reveal interesting features of the circulation. It was clearly apparent that a remote, accurate, relatively inexpensive, long-term tracking system was needed. This was soon provided in the 1970s by satellite


DRIFTERS AND FLOATS

tracking, which revolutionized the tracking of drifters in the ocean. The first satellite tracking of drifters occurred in 1970 using the Interrogation Recording and Location System (IRLS) system flown on Nimbus 3 and 4 satellites. This system measured the slant range and bearing of a radio transmitter on a drifter. The IRLS drifters were very expensive, too expensive at $50 000 to be used in large numbers (but cheap compared to the cost of the satellite). During 1972–73, several drifters were tracked with the (Corporative Application Satellite) EOLE system, which incorporated Doppler measurements of the drifter radio transmissions to determine position. In the mid-1970s, NASA developed the Random Access Measurement System (RAMS), which used Doppler measurements and was flown on the Nimbus 6 satellite. The radio transmitters were relatively inexpensive at $1300, and tracking was provided free by NASA (as proof of concept), which enabled many oceanographers to begin satellite tracking of surface drifters. The modern Doppler-based satellite tracking used today, the French Argos system flown on polar orbiting satellites, is similar to the early RAMS but provides improved position performance. The modern drifter transmitter emits a 0.5-W signal at 402 MHz approximately every minute. Positions are obtained by Service Argos about six times per day in the equatorial region increasing to 15 times per day near 601 N. Position errors are around 300 m. The cost per day of satellite tracking is around $10, which becomes quite expensive for continuous yearlong trajectories. To reduce costs, some drifters have been programmed to transmit only one day out of three or for one-third of each day. This causes gaps in the trajectories that need to be interpolated. Recently, drifters with Global Positioning System (GPS) receivers have been deployed to obtain more accurate (B10 m) and virtually continuous fixes. GPS position and sensor data need to be transmitted to shore via the Argos system or another satellite that can relay data. Experiments are underway using new satellite systems to relay information both ways – to the drifters and to the shore – in order to increase bandwidth, to decrease costs, and to modify sampling. WOCE drifter The development of satellite tracking in the 1970s quickly revealed the weakness of available drifters – most performed poorly and most did not survive long at sea. Many problems needed to be overcome: fishbite, chafe, shockloading, biofouling, corrosion, etc. Early drogues tended to fall off fairly quickly and, since drogue sensors were not used or did not work well, no one knew how

173

long drogues remained attached. Over the years many people tried various approaches to solve these problems, but it was mainly due to the impetus of two large experiments, Tropical Ocean and Global Atmosphere (TOGA) and World Ocean Circulation Experiment (WOCE), and with the persistent efforts of Peter Niiler and colleagues that a good surface drifter was finally developed, standardized, and deployed in large numbers. The so-called WOCE drifters have good water-following characteristics and the slip of the drogue has been calibrated in different conditions. The WOCE drifter works fairly reliably and often survives longer than a year at sea. As of March 2007 there were around 1300 drifters being tracked in the oceans as part of the Global Drifter Program. Data assembly and quality control is performed by the Drifter Data Assembly Center at the National Oceanic and Atmospheric Administration (NOAA) in Miami, Florida. Recent analyses include mapping surface velocity over broad regions and the generation of maps of mean sea level pressure based on drifter measurements. The WOCE drifter consists of a spherical surface float 35 cm in diameter, a 0.56-cm diameter plastic-impregnated wire tether, with a 20-cm diameter subsurface float located at 275 cm below the surface and a drogue in the shape of a 644-cmlong cloth cylinder 92 cm in diameter with circular holes in its sides (Figure 2). The fiberglass surface float contains a radio transmitter, batteries, antenna, and sensors including a thermometer and a submergence sensor that indicates if the drogue is attached. Additional sensors can be added to measure conductivity, atmospheric pressure, light, sound, etc. The basic WOCE drifter costs around $2500, ready for deployment. The WOCE drifter’s drogue is centered at a depth of 15 m below the sea surface. The ratio of the drag area of the drogue to the drag area of tether and float is around 41:1, which results in the drogue’s slip through the water being less than 1 cm s 1 in winds of 10 m s 1. The slip was measured to be proportional to wind speed and inversely proportional to the drag area ratio. From this information, the slip can be estimated and subtracted from the drifter velocity. Drogue Depth

Drogues have been placed at many different depths to suit particular experiments. The drogues of Coastal Ocean Dynamics Experiment (CODE) drifters developed by Russ Davis were located in the upper meter of the water column to measure the surface velocity. WOCE drogues are placed at 15 m


174

DRIFTERS AND FLOATS

often to track subsurface coherent eddies such as meddies. ∅ 35 cm

307 cm

Subsurface Floats Many kinds of freely drifting subsurface floats are being used to measure ocean currents, although most floats are usually either autonomous or acoustic. An autonomous float measures a series of subsurface displacements and velocities between periodic surface satellite fixes. An acoustic float measures continuous subsurface trajectories and velocities using acoustic tracking. Acoustic floats provide high-resolution ocean trajectories but require an acoustic tracking array and the effort to calculate subsurface positions, both of which add cost. Thousands of autonomous and acoustic floats have been deployed to measure the general circulation in the world ocean at various depths but concentrated near 800 m. Historical and WOCE era float data can be seen and obtained on the WOCE float website along with references to a series of detailed scientific papers, and newer float data on the Argo website.

∅ 20 cm

1500 cm

∅ 46 cm

WOCE Autonomous Float 644 cm

Drag area ratio = 41.3 92 cm Figure 2 Schematic of WOCE surface drifter. Adapted from Sybrandy AL and Niiler PP (1991) WOCE/TOGA Lagrangian Drifter Construction Manual, SIO Reference 91/6, WOCE Report No. 63. La Jolla, CA: Scripps Institution of Oceanography.

to measure a representative velocity in the Ekman layer but below the fastest surface currents. Many scientists have deployed drogues at around 100 m to measure the geostrophic velocity below the Ekman layer. An argument for the 100-m depth is that it is better to place the drogue below the complicated velocity structures in the Ekman layer, Langmuir circulations, and near-surface convergence regions. An argument against the 100-m depth is that the drag of the long tether and surface float in the relatively fast Ekman layer could create excessive slip of the drogue and bias the drifter measurement of geostrophic velocity. The controversy continues. The 15-m depth is widely used today, but many earlier drifters had deeper drogues at around 100 m. Some drogues have been placed as deep as 500–1000 m

The autonomous WOCE float was developed in the 1990s by Russ Davis and Doug Webb. The float typically drifts submerged for a few weeks at a time and periodically rises to the sea surface where it transmits data and is positioned by the Argos satellite system. After around a day drifting on the surface, the float resubmerges to its mission depth, typically somewhere in the upper kilometer of the ocean, and continues to drift for another few weeks. Around 100 round trips are possible over a lifetime up to 6 years. The float consists of an aluminum pressure hull 1 m in length and 0.17 m in diameter (Figure 3). A hydraulic pump moves oil between internal and external bladders, forcing changes of volume and buoyancy and enabling the float to ascend and descend. An antenna transmits to Argos and a damping plate keeps the float from submerging while it is floating in waves on the sea surface. Some floats are drogued to follow a pressure surface; others can be programmed with active ballasting to follow a particular temperature or density surface; more complicated sampling schemes are possible. Autonomous floats have been equipped with temperature and conductivity sensors to measure vertical profiles as the floats rise to the surface. Electric potential sensors have been added to some floats by Tom Sanford and Doug Webb in order to measure vertical profiles of horizontal velocity. These


DRIFTERS AND FLOATS

175

Antenna port

Evacuation port

Damping disk Argos antenna Internal reservior 107 cm

Argos transmitter Controller and circuit boards

70 cm

Microprocessor battery packs Pump battery packs

Motor Filter Hydraulic pump Latching valve Pressure case

External bladder 17 cm Figure 3 Schematic of Autonomous Lagrangian Circulation Explorer (ALACE) float. For ascent, the hydraulic pump moves oil down from an internal reservior to an external bladder. For descent, the latching valve is opened, allowing oil to flow back into the internal reservior. The antenna shown at the right is mounted on the top hemispherical end cap. Adapted from Davis RE, Webb DC, Regier LA, and Dufour J (1992) The Autonomous Lagrangian Circulation Explorer (ALACE). Journal of Atmospheric and Oceanic Technology 9: 264–285.

floats were recently used to measure the ocean response of hurricanes. Starting in 2000, an array of profiling floats began to be launched as part of an international program called Argo. Plans are to build up the float array reaching 3000 profiling floats by 2007 and to replace them as they are lost. As of March 2007, there are around 2800 Argo floats operational. The floats profile temperature and salinity to a depth of 2000 m and measure velocity at the drift depth near 1000 m. Profiles of temperature and salinity are being used to map large areas of the ocean including velocity at the drift depth and are being incorporated into predictive numerical models. The profiles are being combined with earlier and sparser hydrographic profiles to document oceanic climate changes.

The basic drift data from an autonomous float are subsurface displacements or velocity vectors between surface satellite fixes or between extrapolated positions at the times of descent and ascent. Errors in position are estimated to be around 3 km. The surface drifts cause gaps in the series of subsurface displacements, so the displacements cannot be connected into a continuous subsurface trajectory. Subsurface displacements are typically measured over several weeks, which attenuates the higher-frequency motions of ocean eddies. The main benefit of these floats is that they can be used to map the low-frequency ocean circulation worldwide relatively inexpensively. The cost of a WOCE profiling autonomous float is around $16 000 (cheaper than a day of an oceangoing ship).


176

DRIFTERS AND FLOATS

A recent development is the addition of small wings plus streamlining that transforms the float into a simple autonomous glider as it ascends and descends. These gliders are self-propelled through the ocean with typical horizontal speeds of 30 cm s 1 while moving vertically. Movable internal ballast is used to bank a glider, forcing it to turn. Gliders can be programmed to return to a specific location to hold position, to execute surveys, and to transit ocean basins along lines. Doug Webb at Webb Research Corporation is equipping some with thermal engines that extract energy from the ocean’s thermal stratification in temperate regions in order to continuously power the glider. Phase changes of a fluid are used to force buoyancy changes. Some gliders incorporate suitable navigation and measure vertical profiles of velocity. Recently, fleets of gliders have been directed from shore to survey the evolving structure of coastal regions. Acoustic Floats

In the mid-1950s, Henry Stommel and John Swallow pioneered the concept and development of freely drifting neutrally buoyant acoustic floats to measure subsurface currents. The method uses acoustics because the ocean is relatively transparent to sound propagation. The deep sound channel centered at a depth around 1000 m enables long-range acoustic propagation. The compressibility of hollow aluminum and glass pressure vessels is less than that of water, so that a float can be ballasted to equilibrate and remain near a particular depth or density. For example, if the float is displaced too deep, it compresses less than water and becomes relatively buoyant, rising back to its equilibrium level, which is consequently stable. Once neutrally buoyant, a float can drift with the currents at that depth for long times. In 1955, Swallow built the first successful floats (since called Swallow floats) and tracked them for a few days by means of hydrophones lowered from a ship. A moored buoy provided a reference point for the ship positioning. The first pressure hulls were made out of surplus aluminum scaffolding tubes; Swallow thinned the walls with caustic soda to adjust compressibility and buoyancy. Although several floats failed, two worked successfully, which led to further experiments. In 1957, Swallow tracked deep floats as they drifted rapidly southward offshore of South Carolina, providing the first convincing proof of a swift, narrow southward flowing deep western boundary current previously predicted by Stommel. A second experiment in 1959 tracked deep Swallow floats in the Sargasso Sea west of Bermuda. Instead

of drifting slowly in a generally northward direction as had been predicted, the floats drifted fast and erratically, providing convincing evidence of eddy motions that were much swifter than long-term mean circulation. Previously, the deep interior flow was considered too sluggish to be measured with moored current meters. The discovery of mesoscale variability or ocean eddies by Swallow and James Crease using floats radically changed the perception of deep currents and spurred the further developments of both floats and current meters. Swallow floats had a short acoustic range and required a nearby ship to track them, which was difficult and expensive. It was quickly realized that much longer trajectories were needed in order to measure the ocean variability and the lowerfrequency circulation. Accomplishing this required a neutrally buoyant float capable of transmitting significantly more acoustic energy and operating unattended for long times at great pressures. Second, access was required to military undersea listening stations, so that the acoustic signals could be routinely recorded and used to track the floats. In the late 1960s, Tom Rossby and Doug Webb successfully developed and tested the sound fixing and ranging (SOFAR) float, named after the SOFAR acoustic channel. SOFAR floats transmit a low-frequency (250 Hz) signal that sounds in air somewhat like a faint boat whistle. The acoustic signal spreads horizontally through the SOFAR channel and can be heard at ranges of roughly 2500 km. The acoustic arrival times measured at fixed listening stations are used to calculate distances to the float and to triangulate its position. The first success with a SOFAR float drift of four months in 1969 led to further developments and the first large deployment of floats in 1973 as part of Mid-Ocean Dynamics Experiment (MODE). Very interesting scientific results using the float data led to many more experiments and wider use of floats. Later improvements included swept-frequency coherent signaling in 1974, active depth control in 1976, higher power for longer range in 1980, and microprocessors and better electronics in 1983. Moored autonomous undersea listening stations were developed in 1980, freeing experiments from military stations and enabling floats to be tracked in the Gulf Stream and other regions for the first time. SOFAR floats are large (B5-m long) and heavy (B430 kg), which makes them difficult to use in large numbers. In 1984, Rossby developed the RAFOS (SOFAR spelled backward) float, a much smaller, cheaper float that listens to moored sound sources and at the end of its mission surfaces and reports back data via satellite. This float made it much easier


DRIFTERS AND FLOATS

and cheaper to conduct larger experiments; this style of float was improved and tracked in large numbers in the North and South Atlantic as part of WOCE. Various float groups have collaborated in tracking floats at different depths and in maintaining moored tracking arrays. WOCE RAFOS Float

The modern acoustic RAFOS float consists of a glass hull 8.5 cm in diameter and 150–200-cm long, enclosing an electronic package, Argos beacon, and temperature and pressure sensors (Figure 4). An acoustic transducer and external drop weight are attached to an aluminum end cap on the bottom. RAFOS floats are capable of operating at depths from just below the sea surface to around 4000 m. Usually several times per day they listen and record the times of arrival of 80 s 250-Hz acoustic signals transmitted 8.5 cm Copper foil antenna Glass pressure housing Displacement measuring scale 190_200 cm

Radio transmitter

Electronics

Battery pack Pressure transducer Acoustic hydrophone Ballast weight Compressee (isopycnal model) Figure 4 Schematic of RAFOS acoustic float. Adapted from Rossby HT, Dorson D, and Fontaine J (1986) The RAFOS system. Journal of Atmospheric and Oceanic Technology 3: 672–679.

177

from an array of moored undersea sound sources. At the end of the mission, a few months to a few years in length, the float drops an external weight, rises to the sea surface, and transmits recorded times of arrival, temperatures, and pressures to the Argos system. The float remains drifting on the surface for roughly a month before all the data are received and relayed ashore by satellite. A typical float costs $4000–5000 and is considered expendable because it is difficult and expensive to retrieve. The times of arrival are used with the known transmit times of sources and the estimated speed of sound to triangulate the float’s position. A drifting RAFOS float closely follows a pressure surface. A compressee consisting of a spring and piston in a cylinder is sometimes suspended below a RAFOS float, so that it matches the compressibility of seawater. If the compressibilities are the same, the float will remain on or close to a constant density surface and more closely follow water parcels. Some floats have active ballasting and can track a column of water by cycling between two density surfaces. To ballast a RAFOS float, it is weighed in air and water, which gives its volume. Its compressibility is measured by weighing the float at different pressures in a water-filled tank. The amount of weight to be added to make the float neutrally buoyant at the target depth (or density) is calculated using the compressibility and thermal expansion of the float and the temperature and density of the water in the tank and at the target depth. Floats usually equilibrate within 50 m of their target depths or density. Some floats combine acoustic tracking with the active buoyancy of the autonomous float, so that the float can periodically surface and relay data to shore at intervals of a few months. This avoids the long wait for multiyear RAFOS floats to surface and avoids the loss of all data should a float fail during its mission. French MARVOR floats developed by Michel Ollitrault report data back every 3 months and typically survive for 5 years. Drogues have been added to neutrally buoyant floats by Eric d’Asaro to enable them to better measure three-dimensional trajectories. Vertical velocities from these floats are especially interesting in the upper ocean and in the deep convective regions like the Labrador Sea in winter. Another technique used to measure vertical water velocity is the addition of tilted vanes attached to the outside of a float. Water moving vertically past the vanes forces the float to spin and this is measured and recorded. At least two moored sound sources are required to position a RAFOS float. Often three or more are used to improve accuracy. The sources transmit an 80-s swept-frequency 250-Hz signal a few times per day


178

DRIFTERS AND FLOATS

for up to 5 years. Sources are similar to the old SOFAR floats and cost around $33 000. Mooring costs of wire rope, flotation, acoustic release, and other recovery aids can double this figure. Recently, louder, more efficient, and more expensive sound sources have increased tracking ranges up to 4000 km. Errors of acoustic positioning are difficult to estimate and vary depending on the size and shape of the tracking array, the accuracy of float and source clocks, how well the speed of sound is known, etc. Estimates of absolute position errors range from a few kilometers up to 10 km (or more). Fix-to-fix relative errors are usually less than this because some errors cancel and others such as clock errors vary slowly in time. Corrections are made for the Doppler shift caused by a float’s movement toward or away from a source. The typical correction amounts to around 1.3 km for a speed of 10 cm s 1. Tides and inertial oscillations add high-frequency noise to positions and velocities, but since a float integrates these motions, it provides an accurate measure of lower-frequency motions.

See also Acoustics, Deep Ocean. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Ocean Circulation.

Further Reading Burns LG (2007) Tracking Trash: Flotsam, Jetsam, and the Science of Ocean Motion, 56pp. Boston, MA: Houghton Mifflin. D’Asaro EA, Farmer DM, Osse JT, and Dairiki GT (2000) A Lagrangian float. Journal of Atmospheric and Oceanic Technology 13: 1230--1246. Davis RE (2005) Intermediate-depth circulation of the Indian and South Pacific Oceans measured by autonomous floats. Journal of Physical Oceanography 35: 683--707. Davis RE, Sherman JT, and Dufour J (2001) Profiling ALACEs and other advances in autonomous subsurface floats. Journal of Atmospheric and Oceanic Technology 18: 982--993. Davis RE, Webb DC, Regier LA, and Dufour J (1992) The Autonomous Lagrangian Circulation Explorer (ALACE).

Journal of Atmospheric and Oceanic Technology 9: 264--285. Gould WJ (2005) From swallow floats to Argo – the development of neutrally buoyant floats. Deep-Sea Research 52: 529--543. Griffa A, Kirwan AD, Mariano AJ, Rossby T, and Ozgokmen TM (eds.) (2007) Lagrangian Analysis and Prediction of Coastal and Ocean Dynamics. Cambridge, MA: Cambridge University Press. Niiler PP, Davis RE, and White HJ (1987) Water following characteristics of a mixed layer drifter. Deep-Sea Research 24: 1867--1881. Pazan SE and Niiler P (2004) New global drifter data set available. EOS, Transactions of the American Geophysical Union 85(2): 17. Richardson PL (1997) Drifting in the wind: Leeway error in shipdrift data. Deep-Sea Research I 44: 1877--1903. Rossby HT, Dorson D, and Fontaine J (1986) The RAFOS system. Journal of Atmospheric and Oceanic Technology 3: 672--679. Sanford TB, Price JF, Webb DC, and Girton JB (2007) Highly resolved observations and simulations of the ocean response to a hurricane. Geophysical Research Letters 34: L13604. Siedler G, Church J, and Gould J (eds.) (2001) Ocean Circulation and Climate, Observing and Modelling the Global Ocean. London: Harcourt. Swallow JC (1955) A neutrally-buoyant float for measuring deep currents. Deep-Sea Research 3: 74--81. Sybrandy AL and Niiler PP (1991) WOCE/TOGA Lagrangian Drifter Construction Manual, SIO Reference 91/6, WOCE Report No. 63. La Jolla, CA: Scripps Institution of Oceanography.

Relevant Websites http://www.meds-sdmm.dfo-mpo.gc.ca – Archived Drifter Data, Integrated Science Data Management, Fisheries and Oceans Canada. http://www.argo.ucsd.edu – Argo Floats, ARGO home page. http://www.argo.net – Argo Floats, ARGO.NET, The International Argo Project Home Page. http://www.aoml.noaa.gov – Global Drifter Program, Atlantic Oceanographic and Meteorological Laboratory, NOAA. http://wfdac.whoi.edu – WOCE Float Data, WOCE Subsurface Float Data Assembly Center.


DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS M. J. Fogarty, Woods Hole, MA, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 774–781, & 2001, Elsevier Ltd.

Introduction The development of sustainable harvesting strategies for exploited marine species addresses a critically important societal problem. Understanding the sources of variability in exploited marine species and separating the effects of human impacts through harvesting from natural variability is essential in devising effective management approaches. Population dynamics is the study of the continuously changing abundance of plants and animals in space and time. For exploited marine species, population dynamics studies provide the foundation for evaluation of their resilience to exploitation, the determination of optimal harvesting strategies, and the specification of the probable outcomes of alternative management actions. In the following, a ‘population’ is defined as a group of interacting and interbreeding individuals of the same species. A closed population is one in which immigration and emigration are negligible; an open population is one in which dispersal processes do affect abundance levels. A ‘stock’ is a management unit defined on the basis of fishery and/or distinct biological characteristics. ‘Recruitment’ is the number of individuals surviving to the age or size of vulnerability to the fishery. A ‘cohort’ is defined as the individuals born in a specified time interval; a cohort born in a particular year is also referred to as a ‘year class’. Other articles in this encyclopedia focus on population dynamics in the context of multispecies assemblages, climate-related factors, and the ecosystem effects of fishing. Here, the primary emphasis will be on the effects of harvesting at the population level with the recognition that a full understanding of human impacts on exploited marine species can only be attained with a more holistic view incorporating the ecosystem perspective. The conceptual basis for the development of models of exploited populations and the information requirements for these models and an example of an application of these tools in a

fishery assessment and management setting is provided.

Production of Marine Populations The relative balance between increases in biomass due to recruitment and individual growth and losses from mortality due to natural causes and fishing defines the dynamics of exploited populations (Figure 1). The change in biomass of a population over time due to variation in these factors is called net production. For an open population, factors affecting immigration and emigration must also be considered in any evaluation of population change. The integration of information on these fundamental biological and ecological processes is a central focus of population dynamics studies. Prediction of the effects of alternative management regulations or changes in the production characteristics of an exploited population depends on a synthesis of fundamental biological and ecological processes in the form of mathematical models of varying degrees of complexity. Recruitment is often the most variable component of production in many marine populations. Recruitment varies in response to changing environmental conditions as fluctuations in food supply, the activity of predators, and physical conditions affect the growth and survival of the early life stages. The relationship between the reproducing adult population in a given season or year and the resulting recruitment is of particular importance. For a

Recruitment

Immigration

Growth

Population biomass Fishing mortality

Emigration

Natural mortality Figure 1 Components of production and dispersal affecting the biomass of exploited marine populations.


179

DYNAMICS OF EXPLOITED MARINE FISH POPULATIONS

renewable resource, it is of course essential that the replenishment of the population through reproduction should not be adversely affected by exploitation. With respect to the other elements of production, the natural mortality component reflects the effects of biological factors such as predation and disease as well as adverse physical conditions. The increase in size and weight as an individual grows is an important contributor to change in biomass of a cohort over time. Finally, dispersal among populations or subpopulations through immigration and emigration can have important consequences for the persistence of a population and its resilience to exploitation. For example, a harvested population that receives members from an adjacent unexploited subpopulation in effect receives a subsidy that can contribute to its persistence under exploitation.

Maximum sustainable yield

Yield

180

Fishing intensity Figure 2 Relationship between yield (surplus production) and fishing intensity (measured as fishing mortality or fishing effort) under a simple conceptual model incorporating densitydependent feedback processes.

Sustainability

Understanding how marine species will respond to exploitation and the appropriate levels of fishing pressure to ensure continued harvest is an essential component of population dynamics studies. More specifically, the extraction of a yield that is optimal in some defined biological or economic sense is a broadly accepted goal in resource management. In order for a long-term sustainable harvest to be possible, the population must have some capacity to compensate for reductions in population biomass through increased recruitment and growth and/or decreased mortality at one or more life history stages. Mechanisms that underlie such compensatory responses include cannibalism and competition for critical resources. If the population exhibits some form of density dependence in critical processes affecting vital rates, different equilibrium levels of population biomass will exist, corresponding to different levels of fishing pressure. Fluctuations in the physical and biological environment and their effects on the population will result in variation about the equilibrium level. For a population governed by density-dependent feedback processes, harvesting can reduce intraspecific competition or other interactions by reducing density and overall abundance. This results in an increase in the overall productivity of the stock. In the unexploited case, the stock is dominated by larger, older individuals; harvesting shifts the population state to one with a higher proportion of younger, faster-growing (and therefore more productive) individuals. The production generated in this way is called surplus production and, in principle, this production can be taken as yield.

The relationship between surplus production and biomass for a species governed by a simple form of compensatory dynamic is dome-shaped, with a peak at some intermediate level of population biomass. In this simple conceptual model, production is expected to be zero when the biomass is at its highest level (the unexploited state) because of intraspecific interactions such as competition or cannibalism. The production–biomass relationship can be readily translated to one linking yield (surplus production) to fishing intensity. Again, a domed-shaped relationship is expected (Figure 2). At relatively low levels of fishing pressure, the yield is lower than the maximum because some correspondingly low fraction of the population is removed. Conversely, at higher levels of fishing pressure, the productivity of the stock is reduced by removing too high a fraction of the population. The point where yield is the highest is the maximum sustainable yield. All of the points on the curve except where yield is zero are considered to be sustainable yield levels. However, as the population is reduced to low levels, its viability is jeopardized, particularly under variable environmental conditions. It is therefore important not only to specify sustainability as a broad goal of fishery management but to consider optimum harvesting policies that also minimize risk to the population.

Assessing Population Status The development of management strategies involves several components including the determination of the current biomass relative to ‘desired’ levels and projections of how alternative management actions



will affect the population in the future. The reconstruction of past and present population levels typically involves a synthesis of information derived from fishery-dependent and fishery-independent sources. Fishery-dependent information includes factors such as the catch removed from the population and either marketed or discarded at sea, the age and/or size composition of the catch, and the amount of fishing effort and its spatial distribution. Fishery-independent information includes studies to determine the relative abundance of fish though the use of scientific surveys aimed at estimating population levels at different life stages. For example, surveys are often conducted to determine abundance of juvenile and adult fish using modified fishing gears, and others to determine the distribution and abundance of the eggs and larvae of marine species. A careful attempt is made to standardize the methods and gear used over time to ensure that the changes measured from one survey to the next reflect true changes in relative abundance. Other special studies such as mark and recapture experiments are also employed to determine population size and mortality rates. If accurate information on catch levels is available, coupled with information on the size or age composition of the catch, estimates can be made of both population size and mortality rates. The number in the catch removed in a specified time interval, if accurately known, provides an initial minimum estimate of the population size in the sea because at least that many individuals had to have been present to account for the catch. If the fraction of the population removed by harvesting and the fraction dying due to natural sources such as disease and predation can be determined, it is possible to derive an estimate of the actual population size. Knowing the age or size composition of the catch is critical in determining the overall mortality rates. By tracking the changes in the numbers of a cohort over time as it progresses through the fishery, it is possible to estimate the survival rates from one age class to the next and therefore to generate estimates of the population size-at-age. If size but not age composition of the catch is known but we do know the growth rates and the time required to grow from one size class to the next, we can also determine the population size and survival rates. In some cases, the catch adjusted by the amount of fishing effort to obtain that catch can be used as an index of relative abundance. The utility of this index depends on the accuracy of the catch statistics and on the validity of the measure of fishing effort. The latter can be complicated by the fact that several different types of gear may be employed in a single

181

fishery and it will necessary to standardize among the gear types. Similarly, fishing vessels of different sizes have differing fishing power characteristics that require adjustment factors. Finally, technological developments such as advanced navigation and mapping systems, satellite imagery of oceanographic conditions and increasingly sophisticated echosounders used to locate concentrations of target species result in continual increases in the realized fishing power of vessels. Most importantly, it is essential to recognize that fishers are not striving to attain unbiased estimates of overall population size but rather are attempting to maximize their catch rates using all of the experience and tools at their disposal. This must be considered in any attempt to use catch-per-unit-effort as an index of abundance. Scientific surveys are an attempt to provide an independent check on information derived from the fishery itself. Such surveys have proven to be invaluable tools in determining the status of marine populations and the communities and ecosystems within which they are embedded. Typically, fisheryindependent surveys are employed to collect information not only on economically important species but all species that can be adequately sampled by the survey gear, thus providing a broader perspective on changes in the system. In addition, key oceanographic and meteorological measurements are routinely made to index changes in the physical environment. The information noted above is integrated into mathematical models that describe, for example, the decay of a cohort over time under losses due to fishing and other sources. The information from fishery-independent sources can be used to calibrate models operating on fishery-derived data sources in an integrated analysis. In turn, the estimates of population size derived from these models and estimation procedures can be used in models designed to assess alternative harvesting strategies. An example of the application of this overall research approach is provided below for the Icelandic cod population. Estimates of cod recruitment at age 3 years based on applications of models to catch-atage information from the fishery and fishery-independent survey information are provided in Figure 3. Estimates of the number of recruits over a 50year period show fluctuations about a relatively stable level. Estimates of the adult population size for this period are also available and show overall declines during this period as exploitation increased. The relationship between adult biomass and the resulting recruitment for Icelandic cod is provided in Figure 4 along with the predicted fit from a simple model for this relationship. Note that there are substantial deviations from the deterministic curve,


182


capacity is essential if a sustainable harvest is to be possible. An illustration of how this information can be combined with other key aspects of the production of cod to predict yield at different levels of fishing pressure is described later.

800

Recruitment (millions)

700 600 500 400 300

Management Targets and Limits to Exploitation

200 100 0 1925 1935 1945 1955 1965 1975 1985 1995 Year

Figure 3 Estimates of recruitment at age 3 years (millions) for the Icelandic cod population over a 50-year period based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.

800

Recruitment (millions)

700 600 500 400 300 200 100 0 0

400

800

1200

1600

2000

3

Adult biomass (10 t) Figure 4 Relationship between recruitment and adult (spawning) biomass of Icelandic cod based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.

reflecting the effects of variable environmental conditions (and variation attributable to estimation errors). Despite this variability, it is evident that this relationship is nonlinear — recruitment does not increase continuously with increasing adult biomass but rather levels off and even declines at quite high levels of spawning population size. This reflects a form of compensatory response in the adult–recruit relationship. For Atlantic cod populations in general, it is known that cannibalism can be an important regulatory factor that limits recruitment at higher population levels. The model used to develop the predicted relationship between adult biomass and recruitment in Figure 4 is in fact one that incorporates the effect of predation by adults on their progeny. Recall that some form of compensatory

Effective management requires a clear specification of goals and objectives to be achieved. Without an appropriate and pre-agreed target for management, the many conflicting and entrenched interests in the fishery management arena cannot be reconciled. The development of biological and economic reference points has played an integral role in fishery management. In the following, the emphasis will be on biologically based targets for management and on the determination of the resilience of the population to exploitation on the basis of models of the dynamics of exploited populations. One important reference point has already been encountered — the maximum sustainable yield (MSY). The corresponding level of fishing pressure resulting in MSY is also of direct interest. Although the concept of MSY has endured a somewhat controversial history, it remains a cornerstone in many national and international fishery policy statements. For example, in the United States, the Magnuson–Stevens Fishery Management Act of 1996 defines the optimum yield as ‘equal to maximum sustainable yield as reduced by economic, social, or other factors.’ This statement includes an important change in the specification of optimum yield relative to the original legislation introduced in 1976 in which optimum yield was defined as ‘equal to maximum sustainable yield as modified by economic, social, or other factors.’ In the more recent version, MSY and the corresponding fishing mortality rate is taken as a limit to exploitation. Other important biological reference points for management are based on consideration of the effect of harvesting on a cohort once it enters or is recruited to the fishery. By tracking the growth and the fishing and natural mortality over the lifespan of the cohort, it is possible to determine the effects of harvesting on yield and the adult biomass as the level of fishing pressure is changed and/or as we modify the age or size of vulnerability to the fishery. The fishing mortality rate at which yield is maximized (denoted Fmax) is one such reference point (assuming a maximum does in fact exist). An alternative reference point that has been widely applied is defined by the point on the yield-per-recruit curve where the rate of



20 18

2.0

Yield per recruit (kg)

1.8 1.6

16

1.4 1.2

14 12

1.0

10

0.8

8

0.6

6

0.4 0.2

2

4 0

0 0

Spawning biomass per recruit (kg)

0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1.0 Fishing mortality

Figure 5 Yield per recruit (peaked curve) and adult (spawning) biomass per recruit (decaying exponential curve) as a function of fishing mortality for Icelandic cod based on stock assessments conducted under the auspices of the International Council for Exploration of the Sea.

600 500 400

3

Yield (10 t)

change in yield is one-tenth of the rate at the origin (denoted F0.1). An advantage of this reference point is that it always exists (unlike Fmax); further, in cases where Fmax does exist, F0.1 is always lower and therefore leads to more conservative management. An example for Icelandic cod is provided in Figure 5; in this case, FmaxB0.35 and F0.1B0.2. The advantage of this overall approach is that it does not require specific information on the incoming recruitment. Rather, it is possible to express the yield on a per-unit-recruitment basis. A disadvantage of this approach if it is taken alone is that it does not provide direct information on how fishing pressure might affect the replenishment of the population through recruitment. To fully evaluate the effects of fishing on the population, it is possible to combine information from the yield and adult biomass per recruit analysis with the data on recruitment as a function of adult population size to generate a complete life-cycle representation. Fishing reduces the adult biomass and it is possible to estimate the predicted recruitment for each level of fishing as a function of the spawning population (see Figure 4 for the case of Icelandic cod). Multiplying the yield per recruit at each of these fishing rates by the predicted recruitment then gives the total yield. This process is illustrated for the Icelandic cod example in Figure 6 where the predicted equilibrium yield is shown as a function of fishing mortality. Superimposed on this curve are the observed (nonequilibrium) yields for Icelandic cod against the estimated fishing mortality rates. The actual catch data reflect variability in the stock–recruitment relationship due to factors not

183

300 200 100 0 0

0.1 0.2 0.3 0.4

0.5 0.6 0.7 0.8

0.9 1.0

Fishing mortality Figure 6 Predicted equilibrium relationship between yield (surplus production) and fishing mortality (solid line) based on an age-structured model for Icelandic cod incorporating information on the adult–recruitment relationship (see Figure 4), and yield and adult biomass per recruit analyses (see Figure 5). Observed (nonequilibrium) yield and estimated fishing mortality rates for Icelandic cod are shown (dots).

included in the model. The maximum yield is predicted to occur at moderate levels of fishing mortality (FB0.35) and the limiting level of fishing mortality is at FB1.0. Although estimated fishing mortality rates for Icelandic cod decreased at the very end of the available time series, it is clear that Icelandic cod has been overexploited and that the effects of excessive fishing have adversely affected yields. The limiting level of fishing mortality beyond which the probability of stock collapse is high can be shown to be directly related to the rate of recruitment at low spawning stock sizes. Species characterized by a high rate of recruitment at low adult population size are more resilient to exploitation than those with a lower rate.

Shifting Environmental States The discussion has concentrated so far on conditions under which the physical and biological environments affecting the population are relatively stable and do not undergo trends or shifts in state. However, persistent shifts in environmental conditions do occur on decadal timescales with important implications for harvested species. It is useful to distinguish environmental variations in physical factors such as temperature and salinity or biological factors such as prey or predator concentrations that occur on relatively short timescales (seasonal to interannual) from those that occur on longer timescales. The distinction between high-frequency variation on seasonal or annual timescales and low-frequency



variation on decadal timescales has important implications for overall levels of productivity of a population and its resilience to exploitation. With low-frequency variation, the potential interaction between changes in the environment and harvesting is of particular concern because persistent shifts in productivity of the population require change in the biological reference points. Exploitation regimes that are sustainable under one set of environmental conditions may not be under another, lower-productivity pattern. An illustration of this in the context of a simple production model is shown in Figure 7, where a change in the basic productivity level of the population under changing environmental conditions is reflected in a change not only in the overall yield levels attainable but also in the level of fishing pressure at which yield is maximized and at which a stock collapse is predicted. Under the lower-productivity regime depicted, the stock collapses at a fishing pressure that is sustainable (although suboptimally) under higher productivity levels. It is clear that a dynamic concept of maximum sustainable yield and other reference points is required that does account for changing conditions in the biological and physical environments experienced by the population. The integration of population dynamics studies with broader ecological investigations and physical oceanographic research is essential if we are to improve understanding of the effects of harvesting on exploited marine species.

Uncertainty, Risk, and the Precautionary Approach Sustained monitoring of the abundance, demographic characteristics, and productivity of widely distributed marine populations entails special challenges. The precision with which it is possible to measure changes in these key variables depends critically on factors such as funding levels and infrastructure available for both fishery-dependent and fishery-independent programs, and on intrinsic characteristics of the populations themselves such as their degree of heterogeneity in space and time. Some of the principal sources of uncertainty in fishery assessments can be attributed to (a) variability (error) in estimates of population size and demographic characteristics, (b) natural variation in production rates and processes, particularly in recruitment, and (c) lack of complete information on broader ecosystem characteristics that affect the species targeted by harvesting and on the direct and indirect effects of harvesting on the ecosystem. The intrinsic variability of marine populations, communities, and ecosystems

Higher-productivity state

Yield

184

Lower-productivity state

Fishing intensity Figure 7 Yield (surplus production) as a function of fishing intensity under two environmental regimes in which the intrinsic rate of increase of the population is affected by changing productivity.

contributes substantially to these components of uncertainty. It is now commonplace to frame issues concerning human health in terms of risk. Considerations of diet, exposure to chemicals, lifestyle choices, etc. affect the probability that an individual will contract certain diseases. Similarly, it is increasingly common in fisheries stock assessments to describe the risk to the population under alternative management scenarios. Although many specifications of risk are possible, an easily understood definition is that risk is the probability that the population will decline and remain below some specified level. Uncertainty in estimates of key parameters and quantities contributes to risk because of the possibility that errors in estimation or in basic model structures may result in overestimates of population size and productivity, inadvertently resulting in overfishing of the resource. The recognition that many populations of exploited marine species are now fully exploited or overexploited has led to an important reevaluation of management policies. In particular, the need for a more precautionary approach to management has been recognized and integrated into a number of national and international management policy statements. Under the precautionary approach, more conservative management is required in situations where higher levels of uncertainty concerning the stock status and production characteristics exist. The burden of proof that harvesting activities were detrimental to marine populations, communities, and ecosystems has historically (and implicitly) rested with scientists and managers. An important element of the precautionary approach is the recognition that a shift in the burden of proof is required to show that the users of a resource are not adversely impacting the productivity of a population or ecosystem.



Summary Questions relating to the stability and resilience of marine species under exploitation involve fundamental ecological considerations such as the role of density-dependent processes in population regulation, the importance of interspecific interactions such as predation and competition, and the role of the physical environment in the production dynamics of the system. Studies of the dynamics of exploited species incorporating these considerations provide an essential framework for quantitative consideration of the effects of alternative harvesting policies on these populations. The importance of fisheries in an economic and social context has led to intensive efforts on a global basis to understand the dynamics of exploited marine species on broad spatial and temporal scales. These studies have become increasingly important as it has become clear that we are near or have exceeded the apparent limits to fishery production from marine systems. Widespread problems such as overcapitalization and excess capacity of fishing fleets and the prevalence open-access fisheries, however, remain substantial impediments to effective management. The extensive information base available for exploited marine species and recent

185

advances in understanding of ocean ecosystem dynamics can provide a strong foundation for improvements in resource management.

See also Ecosystem Effects of Fishing. Fisheries: Multispecies Dynamics. Fishery Management. Fishery Management, Human Dimension. Marine Fishery Resources, Global State of.

Further Reading Gulland JA (ed.) (1977) Fish Population Dynamics. New York: Wiley Interscience. Gulland JA (1983) Fish Stock Assessment. Gulland JA (ed.) (1988) Fish Population Dynamics, 2nd edn. New York: Wiley Interscience. Hilborn R and Walters CJ (1992) Quantitative Fisheries Stock Assessment. New York: Chapman and Hall. Quinn TJ and Deriso R Quantitative Fish Dynamics. New York: Oxford University Press. Rothschild BJ (1986) Dynamics of Marine Fish Populations. Cambridge, MA: Harvard University Press.


EAST AUSTRALIAN CURRENT G. Cresswell, CSIRO Marine Research, Tasmania, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 783–792, & 2001, Elsevier Ltd.

flow in a narrow band against the western boundary is consequently very strong. At the same time, the westward propagation of Rossby waves accumulates energy against eastern Australia. As a result there is a band several hundred kilometers wide off central eastern Australia where the subsurface water

The East Australian Current The East Australian Current (EAC) is a strong western boundary current akin to the Gulf Stream, the Brazil and the Agulhas Currents, and the Kuroshio. Its maximum speed exceeds 2 m s 1 near the surface and its effects extend down several thousand meters. It has a distinct surface core of warm water with a trough-shaped cross-section that is about 100 km wide and 100 m deep. There is a northward undercurrent beneath the EAC that extends from the upper continental slope down to the abyssal plane at 4500 m. The EAC plays a pivotal role in the life cycles of marine fauna and flora of eastern Australia and weather systems respond to the patterns of warm water that it creates. The strength of the EAC came as a surprise to explorer Captain James Cook: ‘‘Winds southerly, a fresh gale’’, he wrote in his log at sunset on 15 May 1770 as he neared Cape Byron. Seeking more searoom for the night, he headed offshore until, ‘‘having increased our soundings to 78 fathoms, we wore and lay with her head in shorey At daylight we were surprised by finding ourselves farther to the southward than we were in the evening, and yet it had blown strong all night’’. The EAC had carried his ship, Endeavour, southward into the strong winds. The EAC has two main sources (Figure 1): One is the Pacific Ocean South Equatorial Current that flows westward into the Coral Sea between the Solomon Islands at 111S and northern New Caledonia at 191S. As this current nears Australia’s Great Barrier Reef it splits at 141–181S, with part going NW to the Gulf of Papua and part going SE as the first ‘tributary’ of the East Australian Current. The other source is the near-surface layer of salty waters of the central Tasman Sea that are linked to the central Pacific Ocean. North of 251S the salty layer slips beneath the fresher and warmer waters of the tropics. Near Australia, as we will see, the salty waters are taken south to Tasmania by the EAC. What is it that drives the EAC? The curl of the wind stress over much of the South Pacific Ocean moves its waters northward. The southward return

10˚

20˚

Australia 30˚

40˚

50˚S 150˚

160˚E

Figure 1 A schematic diagram of the East Australian Current showing its inflows and the eddies that are associated with it. The current off western Tasmania has been named the Zeehan Current.


187

188

EAST AUSTRALIAN CURRENT

Topex/Poseidon Sea Surface Height-25-DEC-1993

16

17

18

19

20

21

22

23

28˚S

33˚S

38˚S

43˚S 147˚E

152˚E

157˚E

Figure 2 An image of sea surface topography prepared from radar altimeter data from the French-US satellite TOPEX/ POSEIDON. The image was formed from satellite passes over a 17-day period centered on 25 December 1993. The passes were 300 km apart on the ground and the measurements along them were effectively 6 km apart. The data are contoured, but unless a pass goes, for example, over the center of an eddy its peak height will be an underestimate. The contour spacing is 10 cm; it is a little over a meter from the highest to the lowest parts of the sea surface. The EAC flowed down the western side of the ridge in this image. On 28 December, yachts in the Sydney to Hobart race encountered a southerly storm that drove waves into the opposing EAC. The waves steepened, the distance between crests decreased, and 67 of the 104 yachts sought safety in nearby ports owing to failures of rigging, hulls, and crew. Image provided by K. Ridgway, CSIRO.

structure of the upper kilometer has been depressed by up to 300 meters. This means that the temperature at any depth in the band is higher, by as much 51C, than that at the same depth in the neighbouring Tasman Sea. It follows that these waters, down to a ‘depth of no motion’, have lower density and, since a low-density water column will be taller than a high density one, the sea surface in the band stands as a ridge about one meter higher than its surroundings

Figure 3 A NOAA satellite sea surface temperature image of the East Australian Current on 18 November 1991. The temperature scale is at the top of the image. There is a 21 21 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. Note the two streams converging at 301S. A concurrent ship section (Figure 4 suggests that the eastern stream was of high salinity while the stream from the Coral Sea was low salinity.

(Figure 2). In other words, it has a greater steric height. Surface currents are controlled by the shape and slope of the sea surface topography in the same way as winds are controlled by atmospheric pressure patterns. The currents follow contours of elevation



189

_100

35.4

_200

35.2

_300

35

0

25

_100

˚C

20

_3

Depth (m)

35.6

x 10

Depth (m)

30˚S section November 1991 0

_200 15 _300 1

0

0

_1

_0.5

_200

ms

Depth (m)

0.5 _100

_1 _1.5

_300 153.5

154

154.5

155

155.5 156 Longitude (deg)

156.5

157

157.5

158

Figure 4 A ship section along 301S out from Australia of salinity, temperature, and north–south velocity component (measured with an acoustic Doppler current profiler, ADCP). Note the warm, low-salinity feature with the trough-shaped section at the edge of the continental shelf. This is the fastest part of the EAC, but the ADCP data show this current to be much broader and deeper and to be carrying mainly higher-salinity west central Pacific water. The current measurements show the reverse (northward) flow of the EAC meander where the subsurface temperature structure has its greatest slope. On this section the maximum southward speed of the EAC was 1.2 m s 1 and the maximum reverse speed was 1.16 m s 1. The innermost station on this section was at the 50 m isobath and suggested that a slope intrusion had upwelled to the surface because the surface and bottom water temperatures were 201C and 171C, respectively. Near and down from the shelf edge these temperatures were encountered at 100 m and 190 m, respectively.

and are strongest where the slopes of the sea surface are steepest. Similarly, deeper currents respond to patterns of subsurface steric height. The low-salinity tributary from the Coral Sea, which is strongest in late summer, flows toward the western side of the ridge at about 251S, where it joins high-salinity inflow from the east (Figures 3 and 4). The resulting EAC flows southward along the western side of the ridge, with higher sea surface elevation on its left (eastern) side. Peaks and saddles along the ridge force the current to meander, accelerate, and decelerate. The western edge of the EAC can spread in across the narrow continental shelf to influence the nearshore waters, while at the same time, as we will discuss later, driving intrusions of

continental slope water onto the shelf. When the EAC reaches the southern end of the ridge, commonly near 331S, it turns anticyclonically (anticlockwise in the southern hemisphere) until it has reversed direction and is running northward several hundred kilometers offshore. Part of it completes a circuit, rejoining the parent EAC, while the remainder meanders eastward along the Tasman Front toward New Zealand. Various estimates have been made of the volume transport in the various components of the EAC system. From the surface to 2000 m depth the South Equatorial Current carries 50 Sverdrups into the Coral Sea. Of this, half flows southward into the EAC, which, at 301S carries 55 Sv. After meandering


190


15

16

17

18

19

152

20

153 _ 34

_ 35

_ 36

_ 37

Figure 5 A NOAA sea surface temperature image of an eddy near Jervis Bay (351S) in early December 1989. There is a 11 11 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. Note the warm band going around the elliptical eddy, the coastal upwelling, and a small cyclonic eddy centered near the intersection of 361S and the shelf edge. (NOAA 11 Tm 45 S 2 Dec. 1989 15172. ^ 1998 CSIRO.)



to the north, 30 Sv recirculates back into the EAC and 25 Sv moves along the Tasman Front. The transport around a large anticyclonic eddy (next section) is about 55 Sv. Recent work suggests that up to 40 Sv can flow northward just south of the separation point of the EAC, perhaps on the western side of a cyclonic eddy.

Eddies South of the ridge are ‘warm-core’ anticyclonic eddies that are 250 km in diameter with edge speeds of

191

over 1 m s 1 (Figures 5 and 6). Just as Captain Cook had been surprised by the EAC, so was Commander J. Lort Stokes on HMS Beagle in July 1838 surprised by the currents in an eddy: ‘‘yfrom Hobarton we carried a strong fair wind to 40 miles east of Jervis Bay when we experienced a current that set us 40 miles S.E. in 24 hours; this was the more extraordinary as we did not feel it before, and scarcely afterwards’’. The subsurface structure in the eddies is depressed by several hundred meters, such that at 400 m their temperature can be 81C warmer than at the same

Currents across an eddy early December 1989

34˚18′S

Jervis Bay

Eddy centre

Batemans Bay

4 knots

36˚30′S 150˚E

153˚30′

Figure 6 Current measurements with a shipboard ADCP around the time of the satellite image in Figure 5. Note that the current speed reaches almost 2 m s 1 on the western side of the elongated eddy.

150˚E

160˚E Coral Sea

30˚S

A

Tasman Sea

A A

August 1976

40˚S

200

40˚S

200

40˚S

Sydney

Lord Howe Is

0

Sydney

Lord Howe Is

200

200

Lord Howe Is

20

30˚S

30˚S

160˚E

Brisbane

Brisbane

Brisbane

Sydney

150˚E 200

160˚E

200

September 1976

200

150˚E 200

October 1976

Figure 7 A cartoon showing how the westward propagation of an undulation on the Tasman Front causes the ridge of high sea surface elevation, around which flows the EAC, to pinch off to form a new anticyclonic eddy. (From Nilsson and Cresswell, 1986.)


192 12


14

16

18

20

22

Figure 8 A NOAA satellite sea surface temperature image of the East Australian Current running southward and then out to sea on 29 September 1991. There is a 21 21 latitude–longitude grid and the 200 m isobath (roughly the edge of the continental shelf edge) is marked as a black line. The white areas are cloud. There are two anti-cyclonic eddies south of this meander. Note the cascade of warm water from the meander to one eddy and then to the next, as well as the small cyclonic eddies produced by current shear at the edges of all three features. (NOAA 11 TMS 45S 29 Sep. 1991 1615z. 1999 ^ CSIRO.)

depth in the surrounding southern Tasman Sea. The primary eddy formation process is linked to the westward propagation of Rossby Waves along the Tasman Front. These move its meanders and

associated thermocline undulations towards Australia, even though the net flow along the Front is eastward. Several times each year a Rossby Wave reaches the ridge in the sea surface against Australia, initially constricting it and then causing it to pinch off a new eddy (Figure 7). These drift slowly southward. The eddies usually become isolated peaks in the sea surface topography that are not encircled by the surface core of the EAC. However, a transient saddle can form between an eddy and the ridge and this allows the warm EAC to ‘reach’ across to the eddy and encircle it (Figure 8). Part of this warm water may spread inward to cover the surface of the eddy, while part may escape the eddy’s influence after encircling it several times. In summer, when the EAC is strongest, the ridge steps its way southward, successively coalescing with eddies, until it reaches the SE corner of the Australian continent at 381S. These eddies retain their identities, with the ridge then consisting of a chain of eddy peaks and saddles. The elongated ridge opens a pathway for the EAC to cascade southward. The strength of this cascade progressively decreases because part of the EAC that encircles each eddy is lost to the northeast. Remnants of the EAC reach southern Tasmania at 431S via smaller eddies and then overshoot by 200 km into the Southern Ocean (Figure 9). Each day in the fishing season fleets of 20–30 Japanese vessels follow parallel paths several kilometers apart across such eddies as they stream their 100 km long lines to catch tuna. When the ridge retracts to the north it can spawn two or three eddies that are either new or rejuvenated in that the introduction of the new warm EAC water has increased their height above the surrounding sea surface. The eddies have lifetimes of over one year. The speeds in them increase from zero at the center to over 1 m s 1 at the perimeter. An eddy disk does not rotate stiffly: the rotation period decreases from 5 days at the perimeter to 1–2 days near the center (Figure 10). The eddies move along complex paths at speeds ranging from near-stationary up to 30 km d 1. The paths followed by their centers include anticlockwise loops about 200 km across that are described in about one month and these can cause eddies to collide with the continental slope. Such a collision distorts a circular eddy into an ellipse, around which the flow appears to conserve angular momentum, giving highest speeds (>1.5 m s 1) near the minor axes and lowest speeds (B1 m s 1) near the major axes. While an eddy is against the continental slope its southward currents can extend in across the shelf. Eddies regain their near-circular shapes once they move out to sea.



12

13

14

15

16

17

18

19

193

20

(A) Figure 9 Winter and summer satellite sea surface temperature images from the NOAA11 satellite. The temperature scale is across the top of the images. Clouds are white. The shelf edge (200 m isobath) is marked as a thin black line. The winter image (A) shows the EAC apparently arrested off eastern Tasmania, while warm water of the Zeehan Current comes down the west coast and part way up the east coast of Tasmania. The summer image (B) shows the EAC to overshoot Tasmania by about 200 km. It entrains Zeehan Current water as it does this.

Not all eddies rotate anticyclonically. Between the ridge and the nearest warm-core eddy to the south – and between warm-core eddies themselves – can be depressions in the sea surface that are cold-core cyclonic eddies. On their inshore (western) sides these drive northward currents of up to 1 m s 1 near the shelf edge. Also, as the EAC separates from the shelf and slope at the southern end of the ridge it often develops instabilities that are carried along its edge. Each starts as a small meander to the west from which a warm plume then reaches back and around a growing cyclonic eddy. The instabilities can form every five days and are spaced at intervals of about 100 km. Similar small eddies form from the shear at the edges of large anticyclonic eddies. The continual

formation of the small cyclonic eddies, each with doming of the water structure in its interior, lifts richer water into the photic zone where it can photosynthesize. Satellite color measurements reveal the widespread effects of this (Figure 11). Occasionally the southward migration of the ridge will force two anti-cyclonic eddies together so that both are affected: they move several times anticlockwise around one another until they coalesce into a large eddy. The process takes several weeks, during which the current speeds in the pair reach 2 m s 1. The product highlights an interesting property of the anticyclonic eddies: Because their waters, like those of the ridge, are warmer down to several hundred meters than the surrounding waters, in


194


10

11

12

13

14

15

16

17

18

(B) Figure 9 Continued

winter they lose heat at the surface more rapidly and they mix down to more than 400 m depth. The nearconstant salinity and temperature in the mixed layer together serve as a unique signature for an eddy once it acquires a summer ‘cap’ owing to insolation and flooding by the EAC. When eddies coalesce, the new eddy has parts of their signature layers one above the other, according to their relative densities.

Effects on the Continental Shelf We have already mentioned that the EAC (and its eddies) can influence the currents on the continental shelf of eastern Australia. The shelf is narrow, with a width of about 25 km, and its depth at mid-shelf is 60–80 m. The edge of the EAC can reach in to drive

southward currents in excess of 0.5 m s 1 at promontories like Indian Head on Fraser Island, Cape Byron, Smoky Cape, Cape Hawke, Sugarloaf Point, and Jervis Bay. This means that it may be more of an influence on the circulation of the Australian shelf than is the Gulf Stream on the 70–120 km wide US eastern shelf with a typical depth of 30 m. Incursions of the EAC onto the shelf quickly overwhelm existing current patterns, replace large parts of the shelf waters, and appear to be a mechanism for driving cold, nutrient-rich intrusions of slope water from 200–300 m depth in towards the coast. The stress by the EAC on the bottom sets up a bottom boundary layer that moves in across the shelf at a slight angle. Near the coast, and notably downstream of headlands (perhaps because of cyclonic motion that they induce), the intrusions may upwell to the surface.



195

can be seen in the images to carry the chlorophyllrich waters along the shelf and well out to sea.

34˚S

35˚ 120

Summary 36˚

130

The near-surface core of the East Australian Current draws water both from the South Equatorial Current via the Coral Sea and from the central Tasman Sea. It flows southward to about 331S, where it separates from the continent and executes a meander to the north. Part recirculates and part proceeds along the Tasman Front toward New Zealand. The meander spawns 250 km diameter anticyclonic eddies several times each year and these migrate southward. Smaller cyclonic eddies are found throughout the region. The separation point occurs near the southern end of a ridge in the sea surface topography. This ridge moves south and north, coalescing with the anticyclonic eddies that are also highs in the surface topography. The East Australian Current cascades southward along and around these structures, ultimately entering the Southern Ocean south of Tasmania each summer. The waters and currents of the East Australian Current and its eddies can extend in across the narrow continental shelf to the shore. Bottom friction can establish a bottom boundary layer that lifts continental slope water onto the shelf to upwell near the coast, particularly when the winds are northerly.

37˚ 1836 34˚S

35˚

36˚

130

37˚

34˚S

1837

35˚

120

36˚

37˚ 130 1831 38˚ 150˚E

Glossary

152˚ 154˚ _ (h) 22 Apr 14 May 1981 (days 112_134)

Figure 10 The paths followed by three satellite-tracked drifters in an eddy that became distorted into an ellipse after it collided with the continental slope of SE Australia. The dots on the tracks indicate daily intervals. Note the short rotation periods for drifter nearest the eddy center. (From Cresswell and Legeckis (1986) Deep-Sea Research.)

This process can be greatly assisted by northerly winds that drive the surface waters offshore in an Ekman layer, to be replaced by the upwelling of the slope intrusions near the shore. The upwelled waters photosynthesize, producing a peak in productivity at about 20–50 m that is a balance between light, nutrients, and grazing by zooplankton. The green waters are known to mariners and their spectacular patterns are clearly evident in color satellite imagery, contrasting with the transparent deep blue of the nutrient-poor EAC (Figure 11). The edge of the EAC

Ekman layer The wind-driven component of transport in the Ekman or surface boundary layer is directed to the left (right) of the mean wind stress in the southern (northern) hemisphere. Rossby waves Rossby waves occur in the atmosphere and the ocean. In the atmosphere they are high and low pressure cells that, while being carried eastward by strong westerly winds, they in fact propagate westward relative to the mean air flow. In the ocean the mean flow is weaker and undulations in the steric height propagate westward as Rossby waves. Steric height and depth of no motion Measurements of temperature, salinity and pressure are needed to calculate steric height, which is the depth difference in the ocean between two surfaces of constant pressure. The deeper of these is often chosen to be a ‘depth of no motion’ which lies on a constant pressure surface and thus the currents are zero. Steric height differences at the ocean surface across the East Australian Current are about 1 m.


196


2

1 0.7 0.5

Chlorophyll a _ mg m 3

0.3 0.2

0.1

Figure 11 An image of chlorophyll concentration in the SW Tasman Sea inferred from color measurements from the SeaWIFS satellite on 22 November 1997. Clouds are white. In the top right the blue color marks the unproductive waters of the EAC. Inshore of those is a coastal upwelling from which two slugs of chlorophyll-rich waters are carried B200 km southward along the continental shelf. There is an anticyclonic eddy at 381S that has high chlorophyll concentrations, which is hard to understand without knowing the eddy’s recent history. Southward from the mainland to southern Tasmania the chlorophyll concentrations are high, perhaps owing to the confluence and mixing of subantarctic and subtropical waters.

See also

Further Reading

East Australian Current. Ekman Transport and Pumping. Mesoscale Eddies. Pacific Ocean Equatorial Currents. Rossby Waves.

Nilsson CS and Cresswell GR (1980) The formation and evolution of East Australian Current warm-core eddies. Progress in Oceanography 9: 133--183. Tomczak M and Godfrey JS (1994) Regional Oceanography: An Introduction. Oxford: Pergamon.


ECONOMICS OF SEA LEVEL RISE R. S. J. Tol, Economic and Social Research Institute, Dublin, Republic of Ireland & 2009 Elsevier Ltd. All rights reserved.

Introduction The economics of sea level rise is part of the larger area of the economics of climate change. Climate economics is concerned with five broad areas:

• • • • •

What are the economic implications of climate change, and how would this be affected by policies to limit climate change? How would and should people, companies, and government adapt and at what cost? What are the economic implications of greenhouse gas emission reduction? How much should emissions be reduced? How can and should emissions be reduced?

The economics of sea level rise is limited to the first two questions, which can be rephrased as follows:

• • •

What are the costs of sea level rise? What are the costs of adaptation to sea level rise? How would and should societies adapt to sea level rise?

These three questions are discussed below. There are a number of processes that cause the level of the sea to rise. Thermal expansion is the dominant cause for the twenty-first century. Although a few degrees of global warming would expand seawater by only a fraction, the ocean is deep. A 0.01% expansion of a column of 5 km of water would give a sea level rise of 50 cm, which is somewhere in the middle of the projections for 2100. Melting of sea ice does not lead to sea level rise – as the ice currently displaces water. Melting of mountain glaciers does contribute to sea level, but these glaciers are too small to have much of an effect. The same is true for more rapid runoff of surface and groundwater. The large ice sheets on Greenland and Antarctica could contribute to sea level rise, but the science is not yet settled. A complete melting of Greenland and West Antarctica would lead to a sea level rise of some 12 m, and East Antarctica holds some 100 m of sea level. Climate change is unlikely to warm East Antarctica above freezing point, and in fact additional snowfall is likely to store more ice on East Antarctica. Greenland and West Antarctica are much warmer, and climate change may

well imply that these ice caps melt in the next 1000 years or so. It may also be, however, that these ice caps are destabilized – which would mean that sea level would rise by 10–12 m in a matter of centuries rather than millennia. Concern about sea level rise is only one of many reasons to reduce greenhouse gas emissions. In fact, it is only a minor reason, as sea level rise responds only with a great delay to changes in emissions. Although the costs of sea level rise may be substantial, the benefits of reduced sea level rise are limited – for the simple fact that sea level rise can be slowed only marginally. However, avoiding a collapse of the Greenland and West Antarctic ice sheets would justify emission reduction – but for the fact that it is not known by how much or even whether emission abatement would reduce the probability of such a collapse. Therefore, the focus here is on the costs of sea level rise and the only policy considered is adaptation to sea level rise.

What Are the Costs of Sea Level Rise? The impacts of sea level rise are manifold, the most prominent being erosion, episodic flooding, permanent flooding, and saltwater intrusion. These impacts occur onshore as well as near shore/on coastal wetlands, and affect natural and human systems. Most of these impacts can be mitigated with adaptation, but some would get worse with adaptation elsewhere. For instance, dikes protect onshore cities, but prevent wetlands from inland migration, leading to greater losses. The direct costs of erosion and permanent flooding equal the amount of land lost times the value of that land. Ocean front property is often highly valuable, but one should not forget that sea level rise will not result in the loss of the ocean front – the ocean front simply moves inland. Therefore, the average land value may be a better approximation of the true cost than the beach property value. The amount of land lost depends primarily on the type of coast. Steep and rocky coasts would see little impact, while soft cliffs may retreat up to a few hundred meters. Deltas and alluvial plains are at a much greater risk. For the world as a whole, land loss for 1 m of sea level rise is a fraction of a percent, even without additional coastal protection. However, the distribution of land loss is very skewed. Some islands, particularly atolls, would disappear altogether – and this may lead to the disappearance of entire nations


197

198

ECONOMICS OF SEA LEVEL RISE

and cultures. The number of people involved is small, though. This is different for deltas, which tend to be densely populated and heavily used because of superb soils and excellent transport. Without additional protection, the deltas of the Ganges–Brahmaputra, Mekong, and Nile could lose a quarter of their area even for relative modest sea level rise. This would force tens of millions of people to migrate, and ruin the economies of the respective countries. Besides permanent inundation, sea level rise would also cause more frequent and more intense episodic flooding. The costs of this are more difficult to estimate. Conceptually, one would use the difference in the expected annual flood damage. In practice, however, floods are infrequent and the stock-at-risk changes rapidly. This makes estimates particularly uncertain. Furthermore, floods are caused by storms, and climate models cannot yet predict the effect of the enhanced greenhouse effect on storms. That said, tropical cyclones kill only about 10 000 people per year worldwide, and economic damages similarly are only a minor fraction of gross national product (GDP). Even a 10-fold increase because of sea level rise, which is unlikely even without additional protection, would not be a major impact at the global scale. Here, as above, the global average is likely to hide substantial regional differences, but there has been too little research to put numbers on this. Sea level rise would cause salt water to intrude in surface and groundwater near the coast. Saltwater intrusion would require desalination of drinking water, or moving the water inlet upstream. The latter is not an option on small islands, which may lose all freshwater resources long before they are submerged by sea level rise. The economic costs of desalination are relatively small, particularly near the coast, but desalination is energy-intensive and produces brine, which may cause local ecological problems. Because of saltwater intrusion, coastal agriculture would suffer a loss of productivity, and may become impossible. Halophytes (salt-tolerant plants) are a lucrative nice market for vegetables, but not one that is constrained by a lack of brackish water. Halophytes’ defense mechanisms against salt work at the expense of overall plant growth. Saltwater intrusion could induce a shift from agriculture to aquaculture, which may be more profitable. Few studies are available for saltwater intrusion. Sea level rise would erode coastal wetlands, particularly if hard structures protect human occupations. Current estimates suggest that a third of all coastal wetlands could be lost for less than 40-cm sea level rise, and up to half for 70 cm. Coastal wetlands provide many services. They are habitats for fish (also as nurseries), shellfish, and birds (including migratory

birds). Wetlands purify water, and protect coast against storms. Wetlands also provide food and recreation. If wetlands get lost, so do these functions – unless scarce resources are reallocated to provide these services. Estimates of the value of wetlands vary between $100 and $10 000 per hectare, depending on the type of wetland, the services it provides, population density, and per capita income. At present, there are about 70 million ha of coastal wetlands, so the economic loss of sea level rise would be measured in billions of dollars per year. Besides the direct economic costs, there are also higher-order effects. A loss of agricultural land would restrict production and drive up the price of food. Some estimates have that a 25-cm sea level rise would increase the price of food by 0.5%. These effects would not be limited to the affected regions, but would be spread from international trade. Australia, for instance, has few direct effects from sea level rise – and may therefore benefit from increased exports to make up for the reduced production elsewhere. Other markets would be affected too, as farmers would buy more fertilizer to produce more on the remaining land, and as workers elsewhere would demand higher salaries to compensate for the higher cost of living. Such spillovers are small in developed economies, because agriculture is only a small part of the economy. They are much larger in developing countries.

What Are the Costs of Adaptation to Sea Level Rise? There are three basic ways to adapt to sea level rise, although in reality mixes and variations will dominate. The two extremes are protection and retreat. In between lies accommodation. Protection entails such things as building dikes and nourishing beaches – essentially, measures are taken to prevent the impact of sea level rise. Retreat implies giving up land and moving people and infrastructure further inland – essentially, the sea is given free range, but people and their things are moved out of harm’s way. Accommodation means coping with the consequences of sea level rise. Examples include placing houses on stilts and purchasing flood insurance. Besides adaptation, there is also failure to (properly) adapt. In most years, this will imply that adaptation costs (see below) are less. In some years, a storm will come to kill people and damage property. The next section has a more extensive discussion on adaptation. Estimating the costs of protection is straightforward. Dike building and beach nourishment are routine operations practised by many engineering



companies around the world. The same holds for accommodation. Estimates have that the total annual cost of coastal protection is less than 0.01% of global GDP. Again, the average hides the extremes. In small island nations, the protection bill may be over 1% of GDP per year and even reach 10%. Note that a sacrifice of 10% is still better than 100% loss without protection. There may be an issue of scale. Gradual sea level rise would imply a modest extension of the ‘wet engineering’ sector, if that exists, or a limited expense for hiring foreign consultants. However, rapid sea level rise would imply (local) shortages of qualified engineers – and this would drive up the costs and drive down the effectiveness of protection and accommodation. If materials are not available locally, they will need to be transported to the coast. This may be a problem in densely populated deltas. Few studies have looked into this, but those that did suggest that a sea level rise of less than 2 m per century would not cause logistical problems. Scale is therefore only an issue in case of ice sheet collapse. Costing retreat is more difficult. The obvious impact is land loss, but this may be partly offset by wetland gains (see above). Retreat also implies relocation. There are no solid estimates of the costs of forced migration. Again, the issue is scale. People and companies move all the time. A well-planned move by a handful of people would barely register. A hasty retreat by a large number would be noticed. If coasts are well protected, the number of forced migrants would be limited to less than 10 000 a year. If coastal protection fails, or is not attempted, over 100 000 people could be displaced by sea level rise each year. Like impacts, adaptation would have economy-wide implications. Dike building would stimulate the construction sector and crowd out other investments and consumption. In most countries, these effects would be hard to notice, but in small island economies this may lead to stagnation of economic growth.

How Would and Should Societies Adapt to Sea Level Rise? Decision analysis of coastal protection goes back to Von Dantzig, one of the founding fathers of operations research. This has also been applied to additional protection for sea level rise. Some studies simply compare the best guess of the costs of dike building against the best guess of the value of land lost if no dikes were built, and decide to protect or not on the basis of a simple cost–benefit ratio. Other studies use more advanced methods that, however, conceptually boil down to cost–benefit analysis as well. These

199

studies invariably conclude that it is economically optimal to protect most of the populated coast against sea level rise. Estimates go up to 85% protection, and 15% abandonment. The reason for these high protection levels are that coastal protection is relatively cheap, and that low-lying deltas are often very densely populated so that the value per hectare is high even if people are poor. However, coastal protection has rarely been based on cost–benefit analysis, and there is no reason to assume that adaptation to sea level rise will be different. One reason is that coastal hazards manifest themselves irregularly and with unpredictable force. Society tends to downplay most risks, and overemphasize a few. The selection of risks shifts over time, in response to events that may or may not be related to the hazards. As a result, coastal protection tends to be neglected for long periods, interspersed with short periods of frantic activity. Decisions are not always rational. There are always ‘solutions’ waiting for a problem, and under pressure from a public that demands rapid and visible action, politicians may select a ‘solution’ that is in fact more appropriate for a different problem. The result is a fairly haphazard coastal protection policy. Some policies make matters worse. Successful past protection creates a sense of safety and hence neglect, and attracts more people and business to the areas deemed safe. Subsidized flood insurance has the same effect. Local authorities may inflate their safety record, relax their building standards and land zoning, and relocate their budget to attract more people. A fundamental problem is that coastal protection is partly a public good. It is much cheaper to build a dike around a ‘community’ than around every single property within that community. However, that does require that there is an authority at the appropriate level. Sometimes, there is no such authority and homeowners are left to fend for themselves. In other cases, the authority for coastal protection sits at provincial or national level, with civil servants who are occupied with other matters. A number of detailed studies have been published on decision making on adaptation to coastal and other hazards. The unfortunate lesson from this work is that every case is unique – extrapolation is not possible except at the bland, conceptual level.

Conclusion The economics of climate change are still at a formative stage, and so are the economics of sea level rise. First estimates have been developed of the order of magnitude of the problem. Sea level rise is not a substantial economic problem at the global or even the


200


continental scale. It is, however, a substantial problem for a number of countries, and a dominant issue for some local economies. The main issue, therefore, is the distribution of the impacts, rather than the impacts themselves. Future estimates are unlikely to narrow down the current uncertainties. The uncertainties are not so much due to a paucity of data and studies, but are intrinsic to the problem. The impacts of sea level rise are local, complex, and in a distant future. The priority for economic research should be in developing dynamic models of economies and their interaction with the coast, to replace the current, static assessments.

See also

Monsoons, History of. Paleoceanography: the Greenhouse World. Salt Marshes and Mud Flats. Sea Level Change.

Further Reading Burton I, Kates RW, and White GF (1993) The Environment as Hazard, 2nd edn. New York: The Guilford Press. Nicholls RJ, Wong PP, Burkett VR, et al. (2007) Coastal systems and low-lying areas. In: Parry ML, Canziani O, Palutikof J, van der Linden P, and Hanson C (eds.) Climate Change 2007: Impacts, Adaptation and Vulnerability – Contribution of Working Group II to the Fourth Assessment Report of the Intergovernmental Panel on Climate Change, pp. 315--356. Cambridge, UK: Cambridge University Press.

Abrupt Climate Change. Beaches, Physical Processes Affecting. Coastal Zone Management.


ECOSYSTEM EFFECTS OF FISHING S. J. Hall, Flinders University, Adelaide, SA, Australia Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 793–799, & 2001, Elsevier Ltd.

Introduction In comparison with conventional fisheries biology, which examines the population dynamics of target stocks, there have been relatively few research programs that consider the wider implications of fishing activity and its effects on ecosystems. With growing recognition of the need to conduct and manage our activities within a wider, more environmentally sensitive framework, however, the effects of fishing on ecosystems is increasingly being debated by scientists and policy makers around the world. As with many other activities such as waste disposal, chemical usage or energy policies, scientists and politicians are being asked whether they fully understand the ecological consequences of fishing activity. The scale of biomass removals and its spatial extent make fishing activity a strong candidate for effecting large-scale change to marine systems. Coarse global scale analyses provide a picture of our fish harvesting activities as being comparable to terrestrial agriculture, when expressed as a proportion of the earth’s productive capacity. It has been estimated that 8% of global aquatic primary production was necessary to support the world’s fish catches in the early 1980s, including a 27 million tonne estimate of discards (see below). Perhaps the most appropriate comparison is with terrestrial systems, where almost 40% of primary productivity is used directly or indirectly by humans. Although 8% for marine systems

may seem a rather moderate figure in the light of terrestrial demands, if one looks on a regional basis, the requirements for upwelling and shelf systems, where we obtain most fisheries resources, are comparable to the terrestrial situation, ranging from 24 to 35% (Table 1). Bearing in mind that the coastal seas are rather less accessible to humans than the land, these values for fisheries seem considerable, leading many to agree that current levels of fishing – and certainly any increases – are likely to result in substantial changes in the ecosystems involved. It is generally accepted that the majority of the world’s fish stocks are fully or overexploited. When considering ecosystem effects it is useful to distinguish between the direct and indirect effects of fishing. Direct effects can be summarized as follows: 1. fishing mortality on species populations, either by catching them (and landing or throwing them back), by killing them during the fishing process without actually retaining them in the gear or by exposing or damaging them and making them vulnerable to scavengers and other predators; 2. increasing the food available to other species in the system by discarding unwanted fish, fish offal and benthos; 3. disturbing and/or destroying habitats by the action of some fishing gears. In contrast, indirect effects concern the knock-on consequences that follow from these direct effects, for example, the changes in the abundances of predators, prey and competitors of fished species that might occur due to the reductions in the abundance of target species caused by fishing, or by the provision of food through discarding of unwanted catch.

Table 1 Global estimates of primary production and the proportion of primary production required to sustain global fish catches in various classes of marine system Ecosystem type

Open ocean Upwellings Tropical shelves Nontropical shelves Coastal reef systems

Area (106 km2)

332.0 0.8 8.6 18.4 2.0

Primary production (g C m 2 y 103 973 310 310 890

Catch (g m 2 y

1

)

Discards (g m 2 y

1

)

Mean % of primary production

95% CI

1.8 25.1 24.2 35.3 8.3

1.3–2.7 17.8–47.9 16.1–48.8 19.2–85.5 5.4–19.8

1

) 0.01 22.2 2.2 1.6 8.0

0.002 3.36 0.671 0.706 2.51

Reproduced from Hall (1999).


201

202

ECOSYSTEM EFFECTS OF FISHING

By-catch and Discards In many areas of the world a wide variety of fishing gears are used, each focusing on one or a few species. Unfortunately, this focus does not mean that nontarget species, sexes or size-classes are excluded from catches. Target catch is usually defined as ‘the catch of a species or species assemblage that is primarily sought in a fishery’ – nontarget catch, or by-catch as it is usually called, is the converse. By-catch can then be further classified as incidental catch, which is not targeted but has commercial value and is likely to be retained if fishing regulations allow it and discard catch, which has no commercial value and is returned to the sea. The problem of by-catch and discarding is probably one of the most important facing the global fishing industry today. The threat to species populations, the wastefulness of the activity and the difficulties undocumented discarding poses for fish stock assessment are all major issues. A recent published estimate of the annual total discards was approximately 27 million tonnes, based on a target catch of 77 million tonnes. This figure, however, did not include by-catch from recreational fisheries, which could add substantially to the total removals, and the estimate is subject to considerable uncertainty. Figure 1 shows how these discard figures break down on a regional basis. Just over one-third of the total discards occur in the Northwest Pacific,

arising from fisheries for crabs, mackerels, Alaskan pollock, cod and shrimp, the latter accounting for about 45% of the total. The second ranked region is the Northeast Atlantic where large whitefish fisheries for haddock, whiting, cod, pout, plaice and other flatfish are the primary sources. Somewhat surprisingly, capelin is also a rather important contributor to the total, primarily because capelin are discarded due to size, condition and other marketrelated factors. The third place in world rankings is the West Central Pacific, arising largely through the action of shrimp fisheries. These fisheries, prosecuted mainly off the Thai, Indonesian and Philippine coasts, accounted for 50% of the total by-catch for the region, although fisheries for scad, crab and tuna are also substantial contributors. Interestingly, the South East Pacific ranks fourth, not because the fisheries in the area have high discard ratios (on the contrary, the ratios for the major anchoveta and pilchard fisheries are only 1–3%), but simply due to the enormous size of the total catch. For the remaining tropical regions, by-catch is again dominated by the actions of shrimp fisheries, although some crab fisheries are also significant. One characteristic difference between temperate and tropical fishery discards is worthy of note. In the tropics, where shrimp fisheries dominate the statistics, discards mainly comprise small-bodied species which mature at under 20 cm and weigh less than 100 g. In contrast, for the temperate and subarctic

9.1 3.7 0.7

0.9

0.6 1.6

2.8

0.6 1.5

0.8 2.6 0.3

0.8

0.01 0.0001

0.03

5 million tonnes

Figure 1 The regional distribution of discards. (Reproduced from Hall, 1999.)


0.3

0.8


regions discards are generally dominated by sublegal and legal sizes of commercially important, largerbodied species. Thus, in the temperate zone discarding is not only an ecological issue, it is also a fisheries management issue in the strictest sense. Fish are being discarded which, if left alone, would form part of the future commercial catch. A cause of particular concern is the incidental catch of larger vertebrate fauna such as turtles, elasmobranchs and marine mammals. Catch rates for these taxa are generally highest in gillnet fisheries, which increased dramatically in the 1970s and 1980s, particularly for salmonids, squid and tuna. In some fisheries, the numbers caught can be very substantial. In the high seas longline, purse seine and driftnet fisheries for tuna and billfish, for example, migratory sharks form a large component of the catch, with some 84 000 tonnes estimated to have been caught from the central and south Pacific in 1989. As with the other nonteleost taxa for which bycatch effects are a concern, the life-history characteristics of sharks make them particularly vulnerable to fishing pressure. Slow growth, late age at maturity, low fecundity and natural mortality, and a close stock recruitment relationship all conspire against these taxa. Such life-history attributes have also led to marked alterations in the absolute and relative abundance of ray species in the North and Irish Sea, which are subject to by-catch mortality from trawl fisheries. In the Irish Sea for example, the ‘common skate’ (Raja batis) is now rarely caught. For some species (e.g., some species of albatross and turtle species) levels of by-catch are so great that populations are under threat. But even if the mortality rates are not this great (or the data are inadequate) there is a legitimate animal welfare perspective which argues for strenuous efforts to limit by-catch mortalities regardless of population effects. Few people like the idea of turtles or dolphins being needlessly drowned in fishing nets, regardless of whether they will become locally or globally extinct if they continue to be caught. Although declines in populations as a result of bycatch are the most obvious effect, there are also examples where populations have increased because of the increase in food supply resulting from discarding. The most notable among these are seabirds in the North Sea, where in one year it was estimated that approximately 55 000 tonnes of offal, 206 000 tonnes of roundfish, 38 000 tonnes of flatfish, 2000 tonnes of elasmobranchs and 9000 tonnes of benthic invertebrates were consumed by seabirds. There is good evidence that populations of scavenging seabirds in the North Sea are substantially larger than they would be without the extra food provided by discards.

203

Solving the By-catch and Discard Problem

There is no universally applicable solution for mitigating by-catch and discard problems. Each fishery has to be examined separately (often with independent observers on fishing vessels) and the relative merits of alternative approaches assessed. One obvious route to reducing unwanted catch, however, is to increase the selectivity of the fishing method in some way. In trawl fisheries, in particular, technical advances, combined with a greater understanding of the behavior of fish in nets has led to the development of new methods to increase selectivity. These methods adopt one of two strategies. The first is to exploit behavioral differences between the various fished species, using devices such as separator trawls, modified ground gear (i.e., the parts of the net that touch the seabed) or modifications to the sweep ropes and bridles that attach to the trawl doors. For example, separator trawls in the Barents Sea have been shown successfully to segregate cod and plaice into a lower net compartment from haddock, which are caught in an upper compartment. In Alaska this approach has been used to allow 40% of bottomassociated halibut to escape while retaining 94% of cod, the target species. The second approach is to exploit the different sizes of species. In many fisheries it is the capture of undersized fish that is the main problem and regulation of minimum permissible mesh size is of course a cornerstone of most fisheries management regimes. Such a measure can often, however, be improved upon. For example, the inclusion of square mesh panels in front of the codend can often allow a greater number of escapees, because the meshes do not close up when the codend becomes full. In addition, recent work that alters the visual stimulus that the net provides by using different colored netting in different parts has been shown to improve the efficiency of such panels considerably. At the other end of the scale excluding large sharks, rays or turtles from the catch can be achieved by fitting solid grids of various kinds. In some fisheries such devices are now mandatory (e.g., turtle exclusion devices in some prawn fisheries), but there is often resistance from fishermen because they can be difficult to handle and catches of target species can fall. For nontrawl fisheries, examples of technical solutions can also often be found. For example, new methods of laying long-lines have been developed to avoid incidental bird capture and dolphin escapement procedures that are now used in high seas purse seine fleets. Although many technical approaches have met with considerable success in different parts of the world it is important to recognize that technical fixes are only part of the solution – the system in which


204


they have to operate must also be considered. The regulations that govern fisheries and the vagaries of the marketplace often create a complex web of incentives and disincentives that drive the discarding practices of fishermen. The situation can be especially complicated in multispecies fisheries.

The Effects of Trawling and Dredging on the Seabed Disturbance of benthic communities by mobile fishing gears is the second major cause for concern over possible ecosystem effects of fishing, threatening nontarget benthic species and perhaps also the longer-term viability of some fisheries themselves if essential fish habitat is being destroyed. With continuing efforts to find unexploited fish resources, hitherto untouched areas are now becoming accessible as new technologies such as chain mats, which protect the belly of the net, are developed. In Australia, for example, new fisheries are developing in deeper water down to depths of 1200 m. A prerequisite for a rational assessment of fishing effects on benthos is an understanding of the distribution, frequency and temporal consistency of bottom trawling. On a global basis, recent estimates obtained using Food and Agriculture Organization catch data from fishing nations suggest that the continental shelves of 75% of the countries of the world which border the sea were exposed to trawling in 1996. It would appear, therefore, that few parts of the world’s continental shelf escape trawling, although it should be borne in mind that in many fisheries trawl effort is highly aggregated. Although we have an appreciation of average conditions, these are derived from a mosaic of patches, some heavily trawled along preferred tows, others avoided by fishermen because they are unprofitable or might damage the gear. Unfortunately, lack of data on the spatial distribution of fishing effort prevents estimates of disturbance at the fine spatial resolution required to obtain a true appreciation of the scale of trawl impacts. Nevertheless, there is little doubt that substantial areas of the world’s continental shelf have been altered by trawling activity. For the most part the responses of benthic communities to trawling and dredging is consistent with the generalized model of how ecologists expect communities to respond, with losses of erect and sessile epifauna, increased dominance by smaller faster-growing species and general reductions in species diversity and evenness. This agreement with the general model is comforting, but we have also learnt that not all communities are equally affected.

For example, it is much more difficult to detect effects in areas where sediments are highly mobile and experience high rates of natural disturbance, whereas boulder or pebble habitats, those supporting rich epifaunal communities that stabilize sediments, reef forming taxa or fauna in habitats experiencing low rates of natural disturbance, seem particularly vulnerable. However, despite the body of experimental data that has examined the impacts of trawling on benthic communities, it is often not possible to deduce the original composition of the fauna in places where experiments have been conducted because data gathered prior to the era of intensive bottomfishing are sparse. This is an important caveat because recent analyses of the few existing historical datasets suggest that larger bodied organisms (both fish and benthos) were more prevalent prior to intensive bottom trawling. Moreover, in general, epifaunal organisms are less prevalent in areas subjected to intensive bottom fishing. Communities dominated by sponges, for example, may take more than a decade to recover, although growth data are notably lacking. Such slow recovery contrasts sharply with habitats such as sand that are restored by physical forces such as tidal currents and wave action. Habitat Modification

An important consequence of trawling and dredging is the reduction in habitat complexity (architecture) that accompanies the removal of sessile epifauna. There is compelling evidence from one tropical system, for example, that loss of structural epibenthos can have important effects on the resident fish community, leading to a shift from a high value community dominated by Lethrinids and Lutjanids to a lower value one dominated by Saurids and Nemipterids. Similar arguments have also been made for temperate systems where structurally rich habitats may support a greater diversity of fish species. Importantly, such effects may not be restricted to the large biotic or abiotic structure provided by large sponges or coral reefs. One could quite imagine, for example, that juveniles of demersal fish on continental shelves might benefit from a high abundance of relatively small physical features (sponges, empty shells, small rocks, etc.) but that over time trawling will gradually lower the physical relief of the habitat with deleterious consequences for some fish species. Such effects may account for notable increases in the dominance of flatfish in both tropical and temperate systems. Our current understanding of the functional role of many of the larger-bodied long-lived species (e.g., as habitat features, bioturbators, etc.) is limited and needs to be addressed to predict the outcome of



permitting chronic fishing disturbance in areas where these animals occur. Although fishing-induced habitat modification is probably most widely caused by mobile gears, it is important to recognize that other fishing methods can also be highly destructive. For coral reef fisheries, dynamite fishing and the use of poisons represent major threats in some parts of the world. Perhaps the only effective approach for mitigating the effects of trawling in vulnerable benthic habitats is to establish marine protected areas in which the activity is prohibited. Given the widespread distribution of trawling, it is not surprising that the establishment of marine protected areas is a key goal for many sectors of the marine conservation movement, although it should be borne in mind that it is not only trawling effects that can be mitigated by the approach. A key driver for the establishment of marine protected areas has come from The World Conservation Union (IUCN) and others who have called for a global representative system of marine protected areas and for national governments to also set up their own systems. A number of nations have already taken such steps, including Australia, Canada and the USA, with other nations likely to follow suit in the future.

Species Interactions Even species that are not directly exploited by a fishery are likely to be affected by the removal of a substantial proportion of their prey, predator or competitor biomass and there are certainly strong indications that interactions with exploited species should be strong enough to lead to population effects elsewhere. For example, an analysis of the energy budgets for six major marine ecosystems found that the major source of mortality for fish is predation by other fish. Predatory interactions may, therefore, be important regulators for marine populations and removing large numbers of target species may lead to knock-on effects. Unfortunately, however, gathering the data necessary to demonstrate such controls is a major task that has rarely been achieved. Without studies directed specifically at the processes underlying the population dynamics of specific groups of species, it is difficult to evaluate the true importance of the effects of fisheries acting through species interactions in marine systems. Despite this caveat, some general effects appear to be emerging. Removing Predators

For communities occupying hard substrata, there is good evidence that some fisheries have reduced

205

predator abundances and that this has led to marked changes lower in the food web. Both temperate hard substrates and coral reefs provide good examples where reductions in predator numbers have led to change in the abundance of prey species that compete for space (e.g., mussels or algae), or in prey that themselves graze on sessile species. Such changes have led in turn to further cascading changes in community composition. For example, in some coral reef systems, removal of predatory fish has led to increases in sea urchin abundance and consequent reductions in coral cover. Examples of strong predator control are much less easy to find in pelagic systems than they are in hard substratum communities. This perhaps suggests that predator control is less important in the pelagos. Alternatively, the lack of evidence may simply reflect our weak powers of observation; it is much harder to get data that would support the predator control theory in the pelagic than it is on a rocky shore. Removing Prey

Fluctuations in the abundance of prey resources can affect a predator’s growth and breeding success. Thus, if prey population collapses are sustained over the longer term due to fishing this will translate into a population decline for the predator. Examples of such effects can be found, particularly for bird species, but also for other taxa such as seals. Since many people have strong emotional attachments to such taxa, there is often intense interest when breeding failures or population declines occur. In the search for a culprit fishing activity is often readily offered as an explanation for the prey decline, or at least as an important contributory factor. In assessing the effect of prey removal, however, one must consider whether the fishery and predator compete for the same portion of the population, either in terms of spatial location or stage in the life cycle. For example, if the predator eats juveniles whose abundance is uncorrelated with the abundance of the fishable stock, the potential for interactions is greatly reduced. Such a feature seems rather common and probably needs to be examined closely in cases where a fishery effect is implicated. Nevertheless, there can be little doubt that unrestrained exploitation increases the likelihood of fisheries collapses and this is turn will take its toll on predator populations. Removing Competitors

Unequivocal demonstrations of competition in most marine systems are rare. Perhaps the only exception to this is for communities occupying hard substrates where competition for space has been demonstrated


206


and can be important in determining community responses to predators (see above). For other systems (e.g., the pelagic or soft-sediment benthos), we can only offer opinions, based on our assessment of the importance of other factors (e.g., predation, low quality food, environmental conditions). One system where fishing activity has been generally accepted to have an impact through competitive effects is the Southern Ocean, where massive reductions in whale populations by past fishing activity has led to apparent increases in the population size, reproduction or growth of taxa such as seals and penguins. A recent assessment, however, has even cast doubt on this interpretation, concluding that there is little evidence that populations have responded to an increase in available resources resulting from a decline in competitor densities. Species Replacements

Despite the difficulties of clearly identifying the ecological mechanism responsible for the changes, there are some examples where fishing is heavily

implicated in large-scale shifts in the species composition of the system and apparent replacement of one group by another. The response of the fish assemblage in the Georges Bank/Gulf of Maine area is, perhaps, the clearest example (Figure 2). During the 1980s the principal groundfish species, flounders and other finfish, declined markedly in abundance after modest increases in the late 1970s. It seems almost certain that the subsequent decline was a direct result of overexploitation by the fishery. In contrast, the elasmobranchs (skates and spiny dogfish) continued to increase during the 1980s. It would appear, therefore, that the elasmobranchs have responded opportunistically to the decline in the other species in the system, perhaps by being able to exploit food resources that were no longer removed by target species. Other possible examples of species replacements are the apparent increase in cephalopod species in the Gulf of Thailand, which coincided with the increase in trawl fishing activity and reduction in the abundance of demersal fish, and the increase in flatfish species that seems to have occurred in the North Sea and elsewhere.

Groundfish and flounders

3.4

Mean weight per tow

20 15 10 5 0

Other finfish

12

Principal pelagics

Mean trophic level

80 60 40 20 0

3.2

3.0

2.8 1950

1970

1960

(A)

8

1980

1990

Year 3.6

4

150

Mean trophic level

0 Skates and spiny dogfish

100 50 0

3.4 1950 3.2

1994 63 66 69 72 75 78 81 84 87 90 93 96 99 Year

Figure 2 Trends in the relative contribution to total biomass (numbers) made by major taxonomic fish groups. Data from National Oceanographic and Atmospheric Administration, National Marine Fisheries Service, Woods Hole, MA, USA (personal communication).

2.8 0 (B)

1

2

3

4

5

Catch (tonnes × 10 6)

Figure 3 (A) Global trends in mean trophic level of fisheries landings from 1950 to 1994. (B) Plot of mean trophic level versus catch for the north-west Atlantic. (Reproduced from Hall, 1999.)



207

Conclusion

See also

A final perspective on the system-level effects of fisheries come from an examination of changes over the last 45 years in the average trophic level at which landed fish were feeding (Figure 3A). This analysis indicates that there has been a decline in mean trophic level from about 3.3 in the early 1950s to 3.1 in 1994. Very large landings of Peruvian anchoveta, which feeds at a low trophic level, account for the marked dip in the time series in the 1960s and early 1970s. When this fishery crashed in 1972–73 the mean trophic level of global landings rose again. For particular regions, where fisheries have been most developed there have been generally consistent declines in trophic level over the last two decades. Plots of mean trophic level against catches give a more revealing insight into the system-level dynamics of fisheries (Figure 3B). Contrary to expectations from simple trophic pyramid arguments, highest catches are not associated with the lowest trophic levels. This is important because it has been suggested in the past that fishing at lower trophic levels will give greater yields because energy losses from transfers up the food chain will be less. It appears, however, that the global trend towards fishing down to the lower trophic levels yields lower catches and generally lower value species – features indicative of fisheries regimes that are badly in need of restoration. Care needs to be taken when interpreting data such as these, particularly because catches of fish at different trophic levels are influenced by a number of factors including the demand for and marketability of taxa and the level of fishing mortality relative to optimum levels. Declines in catches at the end of the time series, for example, may well reflect depleted stocks of fish at all trophic levels. Nevertheless, these analyses are clear warning signs that global fisheries are operating at levels that are certainly inefficient and probably beyond those that are prudent if we wish to prevent continuing change in the trophic structure of marine ecosystems.

Benthic Organisms Overview. Coral Reef and Other Tropical Fisheries. Dynamics of Exploited Marine Fish Populations. Fish Predation and Mortality. Fisheries: Multispecies Dynamics. Fisheries Overview. Large Marine Ecosystems. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Network Analysis of Food Webs. Sea Turtles. Seabirds and Fisheries Interactions.

Further Reading Alverson DL, Freeberg MH, Murawski SA, and Pope JG (1994) A global assessment of bycatch and discards. FAO Fisheries Technical Paper 339, 233pp, Rome. FAO (1996) Precautionary approach to fisheries. Part 1: Guidelines on the precautionary approach to capture fisheries and species introductions. FAO Fisheries Technical Paper 350/1, Rome. FAO (1997) Review of the state of the world fishery resources: marine fisheries. FAO Fisheries Circular no. 920. Rome. Hall SJ (1999) The Effects of Fishing on Marine Ecosystems and Communities. Oxford: Blackwell Science. Jennings S and Kaiser MJ (1998) The effects of fishing on marine ecosystems. Advances in Marine Biology 34: 201--352. Kaiser MJ and deGroot SJ (2000) The Effects of Fishing on Non-target Species and Habitats: Biological, Conservation and Socio-economic Issues. Oxford: Blackwell Science. Pauly D and Christensen V (1995) Primary production required to sustain global fisheries. Nature 374: 255--257. Pauly D, Christensen V, Dalsgaard J, Forese R, and Torres F (1998) Fishing down marine food webs. Science 279: 860--863.


EELS J. D. McCleave, University of Maine, Orono, ME, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 800–809, & 2001, Elsevier Ltd.

Introduction Migratory, catadromous eels of the genus Anguilla have among the most fascinating life cycles of all fishes. Catadromous fishes breed in the ocean but feed and grow in fresh waters or estuaries. This cycle includes a larval or juvenile migration from breeding to feeding area and an adult migration back. These two migrations have been termed denatant and contranatant to imply that the larval migration is accomplished largely by drift with currents and the adult migration by swimming against the currents. These terms get the essence of eel migration, but they fail to capture the complexity and mystery of eel migrations. Anguillid eels are sexually, ecologically, and behaviorally highly adaptive. They occur naturally in a greater diversity of habitats than any other fishes. These statements apply mainly to the feeding and growth stages in continental waters. In contrast, successful spawning and larval survival seems to depend on rather specific oceanic conditions for all Anguilla. Juveniles and adults of Anguilla are elongate, rather cylindrical, darkly pigmented fishes, reaching about 30 cm to nearly 200 cm at maturity depending on species and gender. They have long continuous median fins extending from the anus around the tail and well forward on the back. The contrast with the larvae, termed leptocephali, is extreme. Leptocephali are laterally compressed and deep bodied. They are nearly transparent, with pigment restricted to the retina of the eye. A major metamorphosis occurs between larva and juvenile. This article considers eels of the Family Anguillidae, which is one of about 22 families of eels. Eels are primitive bony fishes. Within the bony fishes (Class Osteichthyes), there are two orders of eels. The Anguilliformes contains 15 families, including spaghetti eels, morays, cutthroat eels, worm eels, snipe eels, and conger eels. The Saccopharyngiformes contains seven families, including deep-sea gulper eels. These two orders are unified by the presence of leptocephali with continuous dorsal, caudal, and anal fins. W. Hulet and R. Robins have argued that the evolution of the leptocephalus, which is in ionic

208

equilibrium with, and nearly isosmotic with sea water, in these primitive fishes was one solution allowing fishes to complete their life cycle in the sea. Only species in the genus Anguilla are truly catadromous, though not all individuals enter fresh water. The other families are marine, ranging from abyssalpelagic to epipelagic to coastal, with juveniles of some species being estuarine. The easily viewed stages in the life cycle of Anguilla occur in fresh water, and they are the only group of eels to enter fresh waters, so they are often called freshwater eels, an unfortunate name, given their extensive migrations at sea.

Taxonomic and Geographic Diversity of Anguilla Fifteen species of Anguilla comprise the Anguillidae (Table 1). These can be grouped into tropical and temperate species on the basis of coastal and freshwater distribution, and of proximity of those distributions in continental waters to the spawning areas. Two temperate species occur in the North Atlantic Ocean, one in the North Pacific, and two in the South Pacific. Eight tropical species are all distributed in the western Pacific and Indian Oceans. Two species extend from the tropics into temperate zones, one in the South Pacific and one in the Indian Ocean. All species require warm, saline, offshore water for successful reproduction. Appropriate currents must be present to transport the larvae toward continental waters. The widely distributed, temperate species use anticyclonic, subtropical gyres for spawning and use associated western boundary currents for distribution of larvae. The European eel is unusual in having its continental distribution on the eastern side of an ocean basin.

Life Cycle Life cycles are best known for the temperate species, but they are undoubtedly similar for the tropical species (Figure 1). Spawning Areas and Times

Spawning of adult eels has never been observed in nature. Spawning areas and times are inferred from the distribution of small leptocephali. A fascinating case was the discovery of the spawning area of the European eel by Danish fishery biologist, J. Schmidt.


EELS

209

Table 1 Taxonomic diversity and continental distribution of the Family Anguillidae with its single genus Anguilla, 15 species and four subspecies Scientific name

English common name

Continental distribution

A. anguilla (Linnaeus, 1758) A. australisb Richardson, 1841

European eel

A. bengalensis bengalensis (Gray, 1831) A. bengalensis labiata (Peters, 1852) A. bicolor bicolor McClelland, 1844 A. bicolor pacifica Schmidt, 1928 A. celebesensis Kaup, 1856 A. dieffenbachii Gray, 1842 A. interioris Whitley, 1938 A. japonica Temminck and Schlegel, 1846 A. malgumora Popta, 1924 A. marmorata Quoy and Gaimard, 1824

Indian mottled eel

(Te)a Iceland, Europe and North Africa from Norway to Morocco, Mediterranean basin to Black Sea, Canary Islands, Azores (Te) Lord Howe Island, east coast of Australia, Tasmania, New Zealand, and Auckland, Chatham, and Norfolk Islands, New Caledonia (Tr) India, Sri Lanka, Myanmar, Andaman Islands, northern Sumatra (Tr) Eastern Africa from Kenya to South Africa

A. megastoma Kaup, 1856 A. mossambica (Peters, 1852) A. obscura Gu¨nther, 1871 A. reinhardtii Steindachner, 1867 A. rostrata (Lesueur, 1817)

Polynesian longfin eel

a b

Shortfin eel

African mottled eel Indonesian shortfin eel — Celebes longfin eel

(Tr) Eastern Africa, Madagascar, India, Myanmar, Sumatra, Java, Timor, north-western Australia (Tr) Eastern Indonesia, New Guinea, Taiwan

New Zealand longfin eel

(Tr) Sumatra, Java, Timor, the Philippines, Celebes, western New Guinea, smaller islands of eastern Indonesia (Te) New Zealand, and Auckland and Chatham Islands

New Guinea eel

(Tr) Eastern New Guinea

Japanese eel

(Te) Northern Vietnam, northern Philippines, Taiwan, China, Korea, and Japan (Tr) Eastern Borneo

— Giant mottled eel

African longfin eel South Pacific eel Speckled longfin eel American eel

(Tr) South Africa, Madagascar, Indonesia, the Philippines, Japan, southern China, Taiwan, eastward through Pacific islands to the Marquesas (Tr) Solomon Islands and New Caledonia eastward to Pitcairn Island (Tr–Te) Eastern Africa from Kenya to South Africa, Madagascar, Mascarenes (Tr) New Guinea to the Society Islands (Te–Tr) Australia from Victoria to Cape York (north tip), New Caledonia, Lord Howe Island (Te) Southern Greenland, eastern North America from Labrador through the Gulf of Mexico and the West Indies to Venezuela and Guyana, Bermuda

Te, temperate species; Tr, tropical species. Australian and New Zealand subspecies are sometimes recognized.

When Schmidt in 1904 caught a 7.5-cm long leptocephalus of the European eel west of the Faroes in the north-eastern Atlantic Ocean, the chase was on. Between 1913 and 1922, Schmidt made four research cruises in the North Atlantic, and commercial vessels collected plankton samples for him. Only in the western North Atlantic were European leptocephali less than 1 cm long captured, in an area 20– 301N and 50–651W. Larger leptocephali were found to the north and east across the Atlantic toward Europe. Schmidt also caught a few small American eel leptocephali west of the locus of small European leptocephali. Recent research by F.-W. Tesch in Germany and at the University of Maine has confirmed that the

south-western portion of the North Atlantic, the Sargasso Sea, is a primary spawning area for both Atlantic Anguilla. From the distribution in space and time of leptocephali less than 1 cm long, the European eel spawns from February through June, primarily March and April, at about 50–751W in a narrow zonal band. The American eel spawns from February through April, at about 53–781W. The two species overlap in space and time. The northern limit to spawning seems to be near-surface frontal zones in the subtropical convergence, which may serve as a cue for the adults to cease migrating and to spawn. Not all areas of the North Atlantic, especially south of 201N, have been adequately sampled to rule out other spawning areas.


210

EELS

Yellow eel Silver eel

Glass eel

Leptocephalus Egg Figure 1 Catadromous life cycle of a temperate species of Anguilla, exemplified here by the American eel, A. rostrata.

Small leptocephali of the Japanese eel have been collected in July in a salinity frontal zone at the northern edge of the North Equatorial Current between 1311 and 1471E, west of the Mariana Islands in the western North Pacific. Spawning areas for the Australian and New Zealand temperate species and for most of the tropical species are speculative. Spawning areas of some tropical species may not be far offshore of the areas of continental distribution. Eggs

In the sea, eggs are presumed to be broadcast into the plankton and fertilized externally. Mature eggs of American, European, and Japanese eels, obtained by hormone injections of females, are transparent and about 1 mm in diameter. Eggs of the Japanese eel hatch in 38–45 h at 231C. Japanese leptocephali are about 2.9 mm long at hatching. Leptocephali

Leptocephali of Anguilla are part of the epipelagic plankton. Their bodies are transparent and laterally compressed, with body height being about 20% of length. A series of W-shaped myomeres (muscle segments) extends from head to tail. Inside the myomeres is an acellular, mucus-like matrix. The myomeres are overlain by a thin epithelium. The head is small (leptocephalus means slender head) and bears a set of fang-like teeth projecting forward. Eyes and olfactory organs are well developed. Dorsal, caudal, and anal fins are continuous around the posterior of the body. Small pectoral fins are present. Leptocephali of

Anguilla sp. are separated from one another primarily by the number of myomeres, e.g. 103–111 for the American eel and 112–119 for the European eel. The mode of nutrition is in debate. Most workers have reported the absence of food in the simple guts of leptocephali, but the ingestion of bacteria, microzooplankton, or gelatinous organisms might go undetected. E. Pfeiler proposed that leptocephali absorb dissolved organic matter from the sea water across their epithelium. N. Mochioka and T. Otake showed that guts of leptocephali other than Anguilla collected at sea do contain larvacean houses, zooplankton fecal pellets, and detrital aggregates (particulate organic matter). Leptocephali may absorb dissolved organic matter across the gut. Larval life lasts from a few months for the tropical species to debatably less than one to more than two years for the European eel. Estimates of larval duration are based on counting putative daily rings in otoliths (ear stones), but otoliths as indicators of age in days have not been validated. Leptocephali of the temperate species grow to about 6–8 cm, whereas those of the tropical species grow to 5–6 cm. Concurrently, leptocephali are gradually transported toward continental waters. Glass Eels

At some time, a dramatic metamorphosis of form and physiology occurs. The body loses height and becomes rounder in cross-section, the larval teeth are resorbed, and the mucus matrix is catabolized. Newly transformed eels, still lacking pigment, are termed glass eels. Initially, they are still pelagic, and they move across the wide or narrow continental shelf into coastal waters. Along the way, they lose their strict pelagic habit and move on and off the bottom. Glass eels entering coastal waters and estuaries may become resident there, or they may continue into fresh waters. The invasion of estuaries and fresh waters is seasonal in the temperate species. At a particular location, the immigration usually peaks over two months, earlier at lower latitudes than at higher latitudes. For the Japanese and American eels and the two New Zealand species, glass eel immigration is in late winter, spring, and early summer. Immigration occurs throughout the year in some tropical species. Elvers and Yellow Eels

Pigmentation develops rapidly after entry into estuarine or fresh waters, and the eels are known as elvers and then yellow eels. Elver refers loosely to small pigmented eels. Yellow eels are named because their ventral surfaces are yellow to white. The dorsal


EELS

surfaces are usually shades of green to brown, plain or mottled depending on species. Yellow eels are in the juvenile growth phase of life. Plasticity and adaptation

Habitat selection Gradual upstream movement of a segment of the yellow eel population occurs for several years, so older yellow eels become distributed from coastal waters to far inland. Eels occur in various habitats including saline coastal waters, estuaries, marshes, large rivers, small streams, lakes, ponds, and even subterranean springs. They occur in highly productive to highly unproductive waters. Temperate species invade waters with near tropical temperatures and waters that are seasonally ice covered. Diet Yellow eels are opportunistic, consuming nearly any live prey that can be captured. Benthic invertebrates predominate in the diets, but fish, including eels, become important to larger eels. Eels respond to local abundances of appropriately sized prey through the seasons. Insect larvae may predominate in early summer and young-of-the-year fishes in late summer. Yellow eels are nocturnal, feeding mostly during the early hours of the night. Sex determination and differentiation Sex is partially or wholly determined by environmental conditions. There are no morphologically differentiated sex chromosomes. Differentiation of the gonads does not occur until the yellow eel phase, with considerable variation in age size at differentiation. For American and European eels, a high population density of small eels seems to result in a high proportion of males and vice versa. Within a river basin, lakes generally have a higher proportion of females than riverine sections, which may reflect a population density or productivity effect. Earlier hypotheses that there was a cline of increasing proportion of female American eels with increasing latitude, and that this was due to longer larval life of females, are not supported by current knowledge. Male-dominated rivers occur in northern as well as southern latitudes, and widely varying sex ratios occur in neighboring rivers. This is indicative of some mechanism for sex determination that acts at the river or habitat level. Growth rate and sexual dimorphism Growth rate varies with length of the growing season and with productivity of the habitat. Growth rate also varies among individuals in the same habitat. For a given age, growth rate is greater for females than males, and females attain greater size at maturity than

211

males. Annual growth rate decreases with age. Because the average age at maturity of females is greater than males, the annual average growth rate to maturity may sometimes be greater for males than females. Determination of age of eels is by counting annual growth rings in the otoliths, but preparation of otoliths has varied among investigators. Accuracy and precision are low, with a tendency to overestimate the age of young eels and underestimate the age of older eels. Therefore, calculated growth rates must be interpreted cautiously. Size and age at maturity Within a species and gender, size is more characteristic than age for when a yellow eel will metamorphose into a silver eel and migrate to sea (Table 2). Studies of ages and sizes at migration of the European eel at 44 sites from Tunisia to Sweden allows generalization. Faster-growing eels of both sexes mature at an earlier age but not a larger size than slower-growing eels. The productivity of the habitat influences growth rate and, therefore, size and age at maturation. Thus, variation in length and age at maturity can occur in different habitats within a restricted geographic range. The length of the growing season and the temperature are negatively correlated with latitude, so age at maturity is strongly correlated with latitude. Both sexes display a time-minimizing strategy, i.e., they mature at the earliest opportunity. However, females mature at a larger size than males because they require sufficient energy to migrate to the spawning area and to produce eggs. Length at migration increases with distance to the spawning area for both sexes. Metamorphosis Toward the end of the yellow phase, many morphological and physiological changes occur, transforming a bottom-oriented yellow eel into an oceanic, pelagic, migratory silver eel. Lipid in the muscle increases to 20–35% of muscle mass in European eels. The gut degenerates, suggesting that silver eels do not feed on migration. The eye increases in diameter by approximately 50%, the number of rod cells in the retina increases, and the spectral absorption maxima shifts more toward blue. This results in increased sensitivity for conditions in the oceanic mesopelagic zone by day and the epipelagic zone by night. The ventral surface of the body becomes whiter and more reflective, increasing the countershading. The swim bladder retial capillaries increase in length from yellow to silver eel, by a factor of 2.5 in American eels, increasing swim-bladder gas deposition rate. Additional guanine is deposited in the swim-bladder


212

EELS

Table 2 Range of mean ages and mean lengths at maturity of silver eels of five species of Anguilla; the European eel is by far the most well studied in this regard Age at maturity

Length at maturity

Species

Sex

Mean age range (y)

Factor a

Nb

Mean length range (cm)

Factor

N

A. anguilla

Male Female Male Female Male Female Male Female Male Female

2.3–15.0 3.4–20.0 3.0–12.7 7.1–19.3 14.2–14.4 19.4–23.6 23.2 34.3–49.4 6.4 8.3

6.5 5.9 4.2 2.7 1.0 1.2 — 1.4 — —

21 28 6 6 3 4 1 2 1 1

31.6–46.0 44.9–86.8 27.7–39.2 41.7–95.7 43.2–46.5 60.9–94.0 62.3–66.6 106.3–115.6 48.3 61.4

1.5 1.9 1.4 1.9 1.1 1.5 1.1 1.1 — —

33 38 9 15 3 4 2 2 1 1

A. rostrata A. australis A. dieffenbachii A. japonica

a b

Factor, largest value divided by smallest value. N, number of geographic locations studied.

Table 3 Fecundity (number of eggs) of females of four species of Anguilla over the size range of eels studied; estimates are probably least accurate for A. anguilla Smaller eels

Larger eels

Species

Length (cm)

Fecundity

Length (cm)

Fecundity

A. A. A. A. A.

65 50 70 45 50

775 410 1 009 1 447 646

85 95 145 115 75

1 3 21 23 2

a b

anguilla australis dieffenbachii rostrataa rostratab

000 000 000 000 000

956 901 374 357 949

000 000 187 000 000

Estimate from Maine, USA, 451N. Estimate from Chesapeake Bay, USA, 371N.

wall, reducing diffusive loss of gas. Premigratory silver American eels maintained swim-bladder inflation at a simulated depth of 150 m compared with 60 m for yellow eels. Silver Eels

Silver eels return to the open ocean, migrate to the spawning area, spawn, and presumably die. The temperate species typically leave fresh and coastal waters in mid–late summer or autumn, earlier at higher latitudes than at lower latitudes. They are presumed to spawn at the next spawning season. However, the journey and spawning have not been witnessed for any species, so the biology of this oceanic stage is speculative. The fecundity of eels increases exponentially with length, ranging from about 0.4 to 25 million eggs depending on species and size (Table 3).

Migrations in the Ocean Silver Eels

In a telemetric study in an estuary, silver American eels migrated seaward by selective tidal stream transport. By ascending into the water column when the tide was ebbing and descending to the bottom when the tide was flooding, eels moved seaward in a saltatory fashion. Directed swimming may also be important in less strongly tidal estuaries. In shallow waters of the North Sea, silver European eels have been shown by telemetry to maintain travel in a given direction regardless of tidal direction, without direct contact with the sea bottom. They also have the ability to move along the tidal axis by selective tidal stream transport. How these mechanisms are used in actual migration is unknown. In deeper waters of the western Mediterranean Sea, silver eels also maintained approximately


EELS

213

Table 4 Estimated lengths of migration of silver eels of three species of Anguilla from various locations to the approximate center of the spawning areasa, assuming travel on a great circle route. Data arranged from south to north. Species

Location

Distance (km)

A. anguilla

Tejo River, Portugal Po River, Italy Loire River, France River Shannon, Ireland IJsselmeer, The Netherlands Lake Vidgan, Sweden Thjo´rsa´ River, Iceland Pearl River, China Shih-Ting River, Taiwan Yangtze River, China Hamana Lake, Japan Naktong-gang River, South Korea Mississippi River, Louisiana, USA Cooper River, South Carolina, USA Chesapeake Bay, Virginia, USA Penobscot Bay, Maine, USA St John’s, Newfoundland, Canada St Lawrence River, Quebec, Canada

4980 8200 5600 4965 7025 8300 5150 2840 2205 2605 2135 2480 2265 1440 1550 2165 2840 3820

A. japonica

A. rostrata

a Assumed spawning locations: A. anguilla 251N 601W; A. japonica 151N 1401E; A. rostrata 251N 681W.

unidirectional movement for hours to days. They also made daily vertical migrations, moving upward at dusk to about 160 m and down at dawn to about 320 m. Routes and rates of silver eel migrations in the open oceans are unknown. I infer from the morphological and physiological changes in the eye, the swim bladder, and the skin that occur at metamorphosis to the silver stage that migration occurs in the epipelagic and upper mesopelagic zones. A model of European eel migration, which combined oriented swimming toward the Sargasso Sea with modeled surface currents, predicted an arrival in the Sargasso Sea somewhat south (15–201N) and east of where the smallest leptocephali have been captured. A second simulated arrival occurred later at about 28–301N. Four migrating females of the New Zealand longfin eel were tagged with archival ‘pop-up’ satellite transmitters, programmed to release after two or three months. All four moved eastward from the New Zealand coast as much as 1000 km. This technology may allow rapid advances in knowledge of silver eel migrations at sea, at least for females of the larger species. Because the spawning areas of many of the species are ill defined, the lengths of migrations of many are unknown. Apparently, many of the tropical species spawn over deep water just off the edge of the

continental shelves, e.g., the Celebes, Molucca, and Banda seas in the western Pacific. In contrast, the migrations of the temperate species are lengthy (Table 4) or presumed so. Leptocephali

Leptocephali of American and European eels o5 mm long are distributed between 50 and 300 m deep both day and night, perhaps indicative of the spawning depth of eels. Larger leptocephali perform daily vertical migrations, which increase in magnitude with increasing body size. Those 5–20 mm long descend from 50–100 m deep by night to 100–150 m deep by day. Those 420 mm long are concentrated at 30–50 m by night and 125–250 m by day. The classical account of the horizontal migration of leptocephali of American, European, and Japanese eels is gradual westward transport south of the subtropical convergences of the Atlantic and Pacific Oceans. The westward transport of the Japanese leptocephali is by the North Equatorial Current. The vertical migration of the leptocephali moves them into the Ekman layer at night, so wind drift influences the trajectories of leptocephali in addition to influence of deeper geostrophic currents. The leptocephali then become entrained in the strong western boundary currents, where they are transported rapidly northward. The western species metamorphose


214

EELS

and detrain from the Gulf Stream and Kuroshio, whereas the European leptocephali, not yet ready to metamorphose, continue eastward in the North Atlantic Current. Some have recently claimed that European leptocephali swim toward Europe by a more direct route from the Sargasso Sea. Their arguments are spatial and temporal. First, leptocephali off the continent of Europe increase in mean length from south to north, suggesting a migration in that direction. However, European leptocephali are found in the Gulf Stream and the North Atlantic Current. There could also be other drift routes eastward, such as an eastward countercurrent associated with the subtropical convergence zone. Secondly, the length of larval life is claimed to be only about 6–7 months on the basis of presumed daily rings in otoliths, versus 2 years or more in the classical account. Whether the migration of European leptocephali is passive or at least partially active cannot currently be resolved. However, back calculation of birth dates on the basis of presumed daily rings of the leptocephali implies that spawning occurs throughout the year. Yet, small leptocephali of the European eel only occur in the Sargasso Sea during part of the year. Unless there is another spawning area and time, the oceanic data are incompatible with the otolith data. Glass Eels

Somewhere, perhaps the edge of the continental shelf, metamorphosis from leptocephalus to glass eel occurs. My speculation is that the diurnal vertical migration of leptocephali brings them into contact with the sea bottom on the shelf, perhaps triggering both metamorphosis and a change in behavior. In near-shore waters and in estuaries, glass eels use selective tidal stream transport to migrate against a net seaward flow. The circatidal vertical migration is probably phased to the local tidal cycle through olfactory cues. Whether the cross-shelf migration is drift on residual currents, selective tidal stream transport, or oriented horizontal swimming is unknown.

Genetics and Panmixia How closely related the eel species are is controversial. Debate has focused primarily on the distinctness of the European and American eels, but applies to all Anguilla. European and American eels separate morphologically on number of vertebrae. The mean vertebral numbers are 114–115 and 107– 108, respectively, with little overlap in ranges. J. Schmidt considered them separate during his extensive research in the North Atlantic in the early 1900s.

In 1959, D. Tucker offered the bold hypotheses that: (1) eels from Europe do not return to breed but die without spawning; (2) the two eels are not separate species but are ecophenotypes of the American eel, their distinguishing characters being environmentally determined during egg and early larval stages; and (3) European populations are maintained by offspring of American eels. A north–south temperature gradient in the Sargasso Sea was proposed as the environmental influence on number of myomeres and vertebrae. Tucker’s hypotheses, though criticized immediately, called attention to the lack of knowledge of the breeding and oceanic biology of eels. Today, Tucker’s hypotheses fail on oceanographic and genetic grounds. The spawning areas of both species overlap considerably in space and time, and they stretch primarily zonally not meridionally. The northern limit of spawning of the two species seems to be the very feature that Tucker invoked to provide the environmental difference. The two are not subject to systematically differing temperature conditions. Small leptocephali captured in single plankton net tows in the spawning area segregate bimodally on myomere numbers. Analyses of nucleotide sequences of mitochondrial DNA (mtDNA) and nuclear DNA from European and American eels show the two to be closely related but distinct species. There are small but consistent differences in cytological characteristics of at least 9 of the 19 pairs of chromosomes. Hybrids of the two species are found in low frequency in Iceland, indicating that genetic isolation is not complete. From Schmidt came the idea that European and American eels were each panmictic, i.e., a single breeding population of each species. Examination of nucleotide sequences of both mitochondrial and nuclear DNA has been used to address the question, with mixed results. European eels with particular sequences collected over wide geographic areas from Morocco and Greece to Sweden and Ireland did or did not cluster geographically in different studies. One study suggested weak structuring of the population, with a southern group (North Africa), a western–northern European group, and an Icelandic group. Another suggested no geographic genetic differentiation. The single study of the American eel using mtDNA showed no geographic structuring among samples collected from the Gulf of Mexico to the Gulf of Maine, 4000 km of coastline. Japanese eels collected at seven sites in Taiwan, two in mainland China, and one in Japan, showed no geographic structuring. MtDNA analysis showed genetic similarity between Anguilla australis from Australia and New Zealand, suggesting they not be treated as subspecies.


EELS

Evolution and Paleoceanography Two studies of evolutionary relationships among nine key species (seven in common) were based on sequence analysis of mtDNA. Both considered that the genus evolved originally in the eastern Indian Ocean, then the Tethys Sea, or Indonesian area, with one suggesting A. celebesensis and one suggesting A. marmorata as the ancestral species A. japonica and A. obscura branched off early. More recently evolved and in the same clade are A. australis, A. mossambica, A. reinhardtii, A. anguilla, and A. rostrata. A. mossambica and A. reinhardtii are sufficiently related that further molecular analysis may suggest a single species. The molecular phylogenies match well with groupings of V. Ege from 1939 based primarily on dentition. The exception is that Ege believed the Japanese eel was closely related to the Atlantic species, a relationship difficult to envision zoogeographically. Anguilla is known from fossils in the early Eocene Epoch, perhaps 50 million years ago (50 Ma), and the family may date from 100 Ma. The evolutionary dispersal of the Austral-Asian species occurred during the time when Australia was moving closer to the IndoPacific islands and when the Indian subcontinent had broken from Africa and was drifting toward Asia. Invasion of the Atlantic by the ancestor of the two Atlantic species probably was through the Tethys Sea and its connection to the Atlantic between Africa and Asia. Closing of the Tethys Sea about 30 Ma isolated them. Separation of the closely related A. australis from the Atlantic ancestor must have occurred prior to the closing. The timing of the separation of A. anguilla and A. rostrata is problematic.

Fisheries and Aquaculture Fisheries occur for glass eels, yellow eels, and silver eels in continental waters. All fisheries are for prereproductive stages. World catches of Anguilla reported to the Food and Agriculture Organization averaged 234 000 t (metric tonnes) in 1995–1998. Asia accounted for about 90% and Europe nearly 7%. Under-reporting of catches is probably widespread. These values are misleading in terms of impact because they combine fisheries for glass eels (a few thousand per kilogram) with fisheries for large silver female eels (a kilogram per eel). Glass eel fisheries are heaviest on the Japanese eel and the European eel, with some commercial harvest of other species. At a peak in the 1970s, glass eel catch in the estuary of the River Loire, France, alone averaged more than 500 t annually. Glass eels are used for human consumption (e.g., in Spain and Portugal), for restocking rivers (e.g., in the Baltic Sea

215

area), and primarily for aquaculture (e.g., in Asia and Europe). Eel culture operations in Asia produce about 120 000 t annually, almost all in China, Taiwan, and Japan. In Europe, eel farms produced about 10 000 t of eels in 1998, mostly in Italy, The Netherlands, and Denmark. Wild and cultured eels are prepared as fresh fish, smoked fish, or kabayaki, traditional Japanese grilled eel with a soy sauce. International trade in live eels and the development of large-scale eel culture has negative as well as positive consequences. European and American eels were accidentally or purposely introduced into Japanese rivers, where in some cases they dominate the eel fauna. In Taiwan, wells drilled to supply water to eel farms resulted in aquifer depletion and land subsidence. The nematode, Anguillicola crassus, which naturally coexists with the Japanese eel, was introduced into the populations of European and American eels. Larval and adult worms infest the swim bladder wall and lumen, with unknown effects on swimming ability of silver eels migrating at sea.

Status of Eel Populations The status of the stocks is best known for Atlantic Anguilla. That of the European eel has been assessed frequently and that of the American eel recently by working groups of the International Council for the Exploration of the Sea. For yellow and silver European eels, fishery-dependent and fishery-independent trends have been largely downward in the last 20–50 years. For American eels, trends have been downward in the last 20 years from peaks in the late 1970s or early 1980s, to very low levels in many cases. It is unknown if those peaks represented unusually high population levels. However, loss of habitat and commercial fishing have contributed to declines. Recruitment of young Atlantic Anguilla from the sea has declined. Long-term data on glass eel recruitment are available in The Netherlands, where the trend was dramatically downward in the 1980s to low levels through the 1990s. This is paralleled by the trend in commercial catch of glass eels in the estuary of the Loire River. The numbers of older yellow American eels moving upstream in the St Lawrence River at an eel ladder declined by three orders of magnitude from the early 1980s to 1999. Two to three order declines in upstream-migrating yellow European eels have occurred in Sweden over the last 50 years. Whether or not recruitment declines are the result of lowered spawning stock is not known. It is possible that regime shifts in North Atlantic oceanic conditions have resulted in decreased survival of leptocephali at


216

EELS

sea or in altered transport pathways for the leptocephali. There are negative correlations between the North Atlantic Oscillation Index, indicative of northern North Atlantic circulation patterns and productivity, and recruitment of glass European eels in The Netherlands (dating to 1938), and recruitment of yellow American eels in the St Lawrence River lagged by four years (dating to 1974).

See also Current Systems in the Atlantic Ocean. Fish Larvae. Fish Migration, Horizontal. Fish Migration, Vertical. Florida Current, Gulf Stream and Labrador Current. Kuroshio and Oyashio Currents. Marine Fishery Resources, Global State of. North Atlantic Oscillation (NAO). Wind Driven Circulation

Glossary Abyssalpelagic Zone of the water column of the ocean encompassing depths between 2000 and 6000 m. Anticyclonic Direction of atmospheric or oceanic circulation around an area of high pressure, clockwise in the Northern Hemisphere and anticlockwise in the Southern Hemisphere. Benthic Organisms dwelling near, on, or in the bottom of the sea. Catadromy Life cycle in which a species of fish breeds in the ocean, migrates into fresh water for growth, and returns to the sea at maturity. Circatidal vertical migration Vertical movement by an organism in phase with the tidal cycle of 12.5 h, with vertical movement occurring during slack tides. Clade A group of organisms, e.g., a group of species, sharing characteristics derived from a common ancestor. Cline Geographic trend in some characteristic of a species or population. Contranatant A migration against the direction of prevailing water current flow. Countershading Color pattern of a pelagic species with a dark, sometimes mottled dorsal surface grading to a silvery ventral surface to reduce the contrast between the animal and its background. Daily (diurnal) vertical migration Vertical movement by an organism in phase with the solar cycle of 24 h, with vertical movement occurring during dusk and dawn, usually upward at dusk and downward at dawn. Denatant A migration along the direction of prevailing water current flow, usually involving drift of young stages.

Detritus Dead organic matter. Ecophenotype The expressed characteristics of a particular subset of a genetically similar population caused by environmental conditions. Ekman layer The near-surface layer, approximately 100 m deep, where wind-generated currents prevail. Epipelagic Zone of the ocean encompassing approximately the upper 200 m. Frontal zone Zone of the ocean in which the gradient (usually horizontal) in features of interest is steep, e.g., temperature, salinity, productivity, fauna. Geostrophic current Currents in balance between a pressure-gradient force (gravity) and the Coriolis deflection. Guanine A double-ringed nitrogenous compound forming part of a nucleotide found in DNA, but here also a crystalline substance deposited in the wall of a fish’s swim bladder to reduce gaseous diffusion. Ionic equilibrium Exhibiting equal concentrations of the same ions inside and outside an organism. Isosmotic Exhibiting equal osmotic pressure inside an outside an organism. Larvacean houses Delicate gelatinous cases secreted, and periodically discarded, by planktonic individuals of the Class Larvacea, Subphylum Urochordata, Phylum Chordata. Leptocephalus The larval stage of eels of the Orders Anguilliformes and Saccopharyngiformes and of tarpons Order Elopiformes and bonefishes and spiny eels Order Albuliformes. Meridional Distribution of oceanic characteristics or organisms on a north–south (longitudinal) axis. Mesopelagic Zone of the water column of the ocean encompassing depths of 200–1000 m. Mitochondrial DNA Genetic material of the mitochondria, the organelles that generate energy for animal cells, which is passed from female to offspring in eggs. Molecular phylogeny Evolutionary relationships among taxonomic groups based on molecular techniques, especially the analysis of nucleotide sequences in DNA or amino acid sequences in proteins. Myomere Muscle segment along the flank of a fish, here a larval eel. North Atlantic Oscillation Index Difference in sea-level atmospheric pressure between Lisbon, Portugal (or Ponta Delgada, Azores) and Reykjavik, Iceland. Nuclear DNA Genetic material of the nucleus of cells, a component of which is passed to offspring from both female and male parents.


EELS

Nucleotide Fundamental structural unit of the nucleic acid group of organic macromolecules, here those involved in information storage (in DNA: units containing adenine, cytosine, guanine or thymine), the sequences of which form codes for protein synthesis. Otolith Concretions of calcium carbonate in a protein matrix deposited in the inner ears of bony fishes, frequently sectioned to examine annual or daily growth. Panmixia The characteristic of a species being composed of a single breeding population. Pelagic Areas of the ocean, or organisms dwelling, well away from the bottom of the sea. Plankton Organisms in the water column of oceans and lakes that have weak swimming abilities, and are wafted by currents. Recruitment The entry of organisms, typically fishes, into the next stage of a life cycle or into a fishery by virtue of growth or migration. Regime shift A transition in oceanic conditions from one quasi-stable state to another. Selective tidal stream transport Mechanism whereby an organism migrates along the tidal axis by ascending into the water column when the tide is flowing in the appropriate direction and descending to hold position near the bottom when the tide is flowing in the inappropriate direction. Subtropical gyre Anticyclonic circular pattern of circulation in each of the major ocean basins between the equator and the temperate zone. Subtropical convergence An area of the ocean where waters of differing characteristics come together, in this case equatorial and temperate waters meeting in the subtropical regions of the Northern and Southern Hemispheres.

217

Swim-bladder retia Countercurrent network of capillaries on the surface of a fish’s swim bladder allowing gas pressure to increase and causing diffusion of gas into the swim bladder. Western boundary current Western geostrophic currents of subtropical circulation gyres which are swift, deep, and narrow. Zonal Distribution of oceanic characteristics or organisms on an east–west (latitudinal) axis.

Further Reading Bruun AF (1963) The breeding of the North Atlantic freshwater-eels. Advances in Marine Biology 1: 137--169. FishBase, A Global Information System on Fishes. http:// www.fishbase.org/search.cfm. Hulet WH and Robins CR (1989) The evolutionary significance of the leptocephalus larva. In: Bo¨hlke EB (ed.) Fishes of the Western North Atlantic, part 9, vol. 2, pp. 669--677. New Haven, CT: Sears Foundation for Marine Research, Yale University. Schmidt J (1922) The breeding places of the eel. Philosophical Transactions of the Royal Society of London, Series B 211: 179--210. Smith DG (1989) Order Anguilliformes, Family Anguillidae, freshwater eels. In: Bo¨hlke EB (ed.) Fishes of the Western North Atlantic, part 9, vol. 1, pp. 25--47. New Haven, CT: Sears Foundation for Marine Research, Yale University. Tesch F-W (1977) The Eel. Biology and Management of Anguillid Eels. London: Chapman and Hall. Tesch F-W (1999) Der Aal, 3rd edn. Berlin: Parey. (In German). Tucker DW (1959) A new solution to the Atlantic eel problem. Nature 183: 495--501. Usui A (1991) Eel Culture. Oxford: Fishing News Books.


EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS I. Boyd and N. Hanson, University of St. Andrews, St. Andrews, UK & 2009 Elsevier Ltd. All rights reserved.

Introduction The eff‘ects of climate change on marine mammals will be caused by changes in the interactions between the physiological state of these animals and the physical changes in their environment caused by climate change. In this article, climate change is defined as a long-term (millennial) trend in the physical climate. This distinguishes it from short-term, regional fluctuations in the physical climate. Marine mammals are warm-blooded vertebrates living in a highly conductive medium often with a steep temperature gradient across the body surface. They also have complex behavioral repertoires that adapt rapidly to changes in the conditions of the external environment. In general, we would expect the changes in the physical environment at the scales envisaged under climate change scenarios to be well within the homeostatic capacity of these species. Effects of climate upon the prey species normally eaten by marine mammals, most of which do not have the same level of homeostatic control to stresses in their physical environment, may be the most likely mechanism of interaction between marine mammals and climate change. However, we should not assume that effects of climate change on marine mammals should necessarily be negative.

Responses to Normal Environmental Variation Marine mammals normally experience variation in their environment that is very large compared with most variance predicted due to climate change. Examples include the temperature gradients that many marine mammals experience while diving through the water column and the extreme patchiness of the prey resources for marine mammals. Marine mammals in the Pacific have life-histories adapted to transient climatic phenomena such as El Ninõ, which oscillate every 4 years or so. Consequently, the morphologies, physiologies, behaviors, and life histories of marine mammals will have evolved to cope with this high level of variance. However, it is generally accepted that

218

climate change is occurring too rapidly for the life histories of marine mammals to adapt to longer periods of adverse conditions than are experienced in examples like El Ninõ. Marine mammals appear to cope with other longer wavelength oscillations including the North Pacific and North Atlantic Oscillations and the Antarctic Circumpolar Wave and it is possible that their life histories have evolved to cope with this type of long wavelength variation. Nevertheless, nonoscillatory climate change could result in nonlinear processes of change in some of the physical and biological features of the environment that are important to some marine mammal species. Although speculative, obvious changes such as the extent of Arctic and Antarctic seasonal ice cover could affect the presence of essential physical habitat for marine mammals as well as food resources, and there may be other changes in the structure of marine mammal habitats that are less obvious and difficult to both identify and quantify. Changes in the trophic structure of the oceans in icebound regions, where the ecology is very reliant on sea ice, may lead the trophic pyramid that supports these top predators to alter substantially. The polar bear is particularly an obvious example of this type of effect where both loss of hunting habitat in the form of sea ice and the potential effects on prey abundance are already having measurable effects upon populations. Similarly, changes in coastal habitats resulting from changes in sea level, changes in run-off and salinity, and changes in nutrient and sediment loads are likely to have important effects on some species of small cetaceans with localized distributions. Many sirenians rely upon seagrass communities and anything that affects the sustainability of this food source is likely to have a negative effect on these species. Many marine mammal species have already experienced range retraction and population depletion because of direct interaction with man. Monk seals appear to be particularly vulnerable because they rely upon small pockets of beach habitat, many of which are threatened directly by man and also by rising sea level. Marine mammals are known to be vulnerable to the effects of toxic algal blooms. Toxins may lead to sublethal effects, such as reduced rates of reproduction as well as direct mortality. Several mortality events including coastal whale and dolphin species as well as seals have been attributed to these effects.


EFFECTS OF CLIMATE CHANGE ON MARINE MAMMALS

Climate change could result in increased frequency of the conditions that lead to such effects, perhaps as a result of interactions between temperature and eutrophication of coastal habitats.

Classifying Effects A common approach to the assessment of the effects of climate change is to divide these into ‘direct’ and ‘indirect’ effects. In this case, ‘direct’ effects are those associated with changes in the physical environment, such as those that affect the availability of suitable habitat. ‘Indirect’ effects are those that operate through the agency of food availability because of changes in ecology, susceptibility to disease, changed exposure to pollution, or changes in competitive interactions. Wu¨rsig et al. in 2002 added a third level of effect which was the result of human activities occurring in response to climate change that tend to increase conflicts between man and marine mammals. This division has little utility in terms of rationalizing the effects of climate change because, in simple terms, the effects will operate ultimately through the availability of suitable habitat. Assessing the effects of climate change rests upon an assessment of whether there is a functional relationship between the availability of suitable habitat and climate and the form of these functional relationships, which will differ between species, has not been determined. The expansion and contraction of suitable habitat can be affected by a broad range of factors and some of these can operate on their own but others are often closely related and synergistic, such as the combined effect of retraction of sea ice upon the availability of breeding habitat for seals and also for the food chains that support these predators.

Evidence for the Effects of Climate Change on Marine Mammals There is no strong evidence that current climate change scenarios are affecting marine mammals although there are studies that suggest some typical effects of climate change could affect marine mammal distribution and abundance. There is an increasing body of literature that links apparent variability in marine mammal abundance, productivity, or behavior with climate change processes. However, with the probable exception of those documenting the changes occurring to the extent of breeding habitat for ringed seals within some sections of the Arctic, and the consequences of this also for polar bears, most of these studies simply reflect a trend toward the interpretation of responses of marine mammals to large-scale regional variability in

219

the physical environment, as has already been well documented in the Pacific for El Ninõ, in terms of climate trends. Long-term trends in the underlying regional ecosystem structure are sometimes extrapolated as evidence of climate change. In few, if any, of these cases is there strong evidence that the physical environmental variability being observed is derived from irreversible trends in climate. Some of the current literature confounds understanding of the responses of marine mammals to regional variability with that of climate change, albeit that an understanding of one may be useful in the interpretation and prediction of the effects of the other. Based upon records of species from strandings, MacLeod et al. in 2005 have suggested that the species diversity of cetaceans around the UK has increased recently and that this may be evidence of range expansion in some species. However, the sample sizes involved are small and there are difficulties in these types of studies accounting for observer effort. This is a common story for marine mammals, and many other marine predators including seabirds, in that, there is a great deal of theory about what the effects of climate change might be but little convincing evidence that backs up these suggestions. Even process studies, involving research on the mechanisms underlying how climate change could affect marine mammals, when considered in detail make a tenuous linkage between the physical variables and the biological response of the marine mammals.

Is Climate Change Research on Marine Mammals Scientific? Although it is beyond dispute that marine mammals respond to physical changes in habitat suitability, the relationship between a particular effect and the response from the marine mammal is seldom clear. Where data from time series are analyzed, as in the case of Forcada et al., they are used to test post hoc for relationships between climate and biological variables. There is a tendency in these circumstances to test for all possible relationships using a range of physical and biological variables. Such post hoc testing is fraught with pitfalls because invariably the final apparently statistically significant relationships are not downweighted in their significance by all the other nonsignificant relationships that were investigated alongside those that proved to be statistically significant. Of course, there may be a priori reasons for accepting that a particular relationship is true, but the approach to examining time series rarely provides an analysis of the relationships that were not statistically significant or the a priori reasons there might be for


220


rejection of these. Consequently, current suggestions from the literature about the potential effects of climate change may be exaggerated because of the strong possibility of the presence of type I and type II statistical error in the assessment process. Moreover, in the great majority of examples, it will be almost impossible to clearly demonstrate effects of climate change, as has been the case with partitioning the variance between a range of causes of the decline of the Steller sea lion (Eumetopias jubatus) in the North Pacific and Bering Sea.

Identifying Situations in Which Climate Change is Likely to Have a Negative Effect on Marine Mammals: Future Work To date, little has been done to build predictive frameworks for assessing the effects of climate change of marine mammals. There have been broad assessments and focused ecological studies but these are a fragile foundation for guiding policy and management, and for identifying populations that are at greatest risk. The resilience of marine mammal populations to climate change will reflect resilience to any other change in habitat quality, that is, it will depend upon the extent of suitable habitat, the degree to which populations currently fill that habitat, the dispersal capacity of the species, and the structure of the current population, including its capacity for increase and demographics. Clearly, populations that are already in a depleted state, or that are dependent upon habitat that is diminishing for reasons other than climate change, will be more vulnerable to the effects of climate change. There are also some, as yet unconvincing, suggestions that habitat degradation may occur through effects of climate upon pollutant burdens. The general demographic characteristics of marine mammal populations are relatively well known so there are simple ways of assessing the risk to populations under different scenarios of demographic stochasticity, population size, and isolation. An analysis of this type could only provide a very broad guide to the types of effects that could be expected but, whereas no such analysis has been carried out to date, this should be seen as a first step in the risk-assessment process. The metapopulation structure of many marine mammal populations will affect resilience to climate change and will be reflected in the dispersal capacity of the population. Again, this type of effect could be included within an analysis of the sensitivity of marine mammal populations to climate change under

different metapopulation structures. A feature of climate change is that it is likely to have global as well as local effects and the sensitivity to the relative contribution from these would be an important feature of such an analysis.

See also Baleen Whales. Marine Mammal Migrations and Movement Patterns. Marine Mammal Overview. Marine Mammal Trophic Levels and Interactions. Marine Mammals: Sperm Whales and Beaked Whales. Sirenians.

Further Reading Atkinson A, Siegel V, Pakhamov E, and Rothery P (2004) Long-term decline in krill stocks and increase in salps within the Southern Ocean. Nature 432: 100--103. Cavalieri DJ, Parkinson CL, and Vinnikov KY (2003) 30year satellite record reveals contrasting Arctic and Antarctic decadal sea ice variability. Geophysical Research Letters 30: 1970 (doi:10.1029/2003GL018031). Derocher E, Lunn N, and Stirling I (2004) Polar bears in a warming climate. Integrative and Comparative Biology 44: 163--176. Ferguson S, Stirling I, and McLoughlin P (2005) Climate change and ringed seal (Phoca hispida) recruitment in western Hudson Bay. Marine Mammal Science 21: 121--135. Forcada J, Trathan P, Reid K, and Murray E (2005) The effects of global climate variability in pup production of Antarctic fur seals. Ecology 86: 2408--2417. Grebmeier J, Overland J, Moore S, et al. (2006) A major ecosystem shift in the northern Bering Sea. Science 311: 1461--1464. Green C and Pershing A (2004) Climate and the conservation biology of North Atlantic right whales: The right whale at the wrong time? Frontiers in Ecology and the Environment 2: 29--34. Heide-Jorgensen MP and Lairde KL (2004) Declining extent of open water refugia for top predators in Baffin Bay and adjacent waters. Ambio 33: 487--494. Hunt G, Stabeno P, Walters G, et al. (2002) Climate change and control of southeastern Bering Sea pelagic ecosystem. Deep Sea Research II 49: 5821--5853. Laidre K and Heide-Jorgensen M (2005) Artic sea ice trends and narwhal vulnerability. Biological Conservation 121: 509--517. Leaper R, Cooke J, Trathan P, Reid K, Rowntree V, and Payne R (2005) Global climate drives southern right whale (Eubaena australis) population dynamics. Biology Letters 2 (doi:10.1098/rsbl.2005.0431). Lusseau RW, Wilson B, Grellier K, Barton TR, Hammond PS, and Thompson PM (2004) Parallel influence of climate on the behaviour of Pacific killer whales and



Atlantic bottlenose dolphins. Ecology Letters 7: 1068--1076. MacDonald R, Harner T, and Fyfe J (2005) Recent climate change in the Artic and its impact on contaminant pathways and interpretation of temporal trend data. Science of the Total Environment 342: 5--86. MacLeod C, Bannon S, Pierce G, et al. (2005) Climate change and the cetacean community of north-west Scotland. Biological Conservation 124: 477--483. McMahon C and Burton C (2005) Climate change and seal survival: Evidence for environmentally mediated changes in elephant seal Mirounga leonina pup survival. Proceedings of the Royal Society B 272: 923--928. Robinson R, Learmouth J, Hutson A, et al. (2005) Climate change and migratory species. BTO Research Report 414. London: Defra. http://www.bto.org/research/reports/ researchrpt_abstracts/2005/RR414%20_summary_ report.pdf (accessed Mar. 2008). Sun L, Liu X, Yin X, Zhu R, Xie Z, and Wang Y (2004) A 1,500-year record of Antarctic seal populations in response to climate change. Polar Biology 27: 495--501.

221

Trillmich F, Ono KA, Costa DP, et al. (1991) The effects of El Ninõ on pinniped populations in the eastern Pacific. In: Trillmich F and Ono KA (eds.) Pinnipeds and El Ninõ: Responses to Environmental Stress, pp. 247--270. Berlin: Springer. Trites A, Miller A, Maschner H, et al. (2006) Bottom up forcing and decline of Stellar Sea Lions in Alaska: Assessing the ocean climate hypothesis. Fisheries Oceanography 16: 46--67. Walther G, Post E, Convey P, et al. (2002) Ecological responses to recent climate change. Nature 416: 389--395. Wu¨rsig B, Reeves RR, and Ortega-Ortiz JG (2002) Global climate change and marine mammals. In: Evans PGH and Raga JA (eds.) Marine Mammals – Biology and Conservation, pp. 589--608. New York: Kluwer Academic/ Plenum Publishers.


EKMAN TRANSPORT AND PUMPING T. K. Chereskin, University of California San Diego, La Jolla, CA, USA J. F. Price, Woods Hole Oceanographic Institution, Woods Hole, Massachusetts, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 809–815, & 2001, Elsevier Ltd.

Introduction Winds blowing along the ocean’s surface exert forces that set the oceans in motion, producing both currents and waves. Separating the wind force into the part that goes into making currents from that which goes into making waves is in fact very difficult. Conceptually, normal forces (i.e., think of the wind beating on the ocean surface like a drum) create waves, and tangential forces (i.e., frictional stresses exerted by the wind pulling on the sea surface) go into making currents. Although there are wind-generated currents that flow in a direction more or less downwind, the currents driven by the steady or slowly varying (compared to the period of the earth’s rotation) wind stress flow in a direction that is quite different from the wind direction, sometimes by more than 901, due to the combined effects of the wind force and the earth’s rotation. These wind-driven currents, commonly called Ekman layer currents in recognition of the Swedish oceanographer W. Ekman who first described their dynamics, are the topic of this article, which has three themes: (1) the local dynamics of Ekman layer currents; (2) the spatial variation of wind stress and the resulting spatial variation of the Ekman layer currents; and (3) the effects of Ekman layer currents on the physical and biological environment of the oceans. The effects of Ekman layer currents are quite far reaching despite the fact that the currents themselves are usually modest, typically no more than 0.05– 0.1 m s1 and smaller than many other currents. The importance of Ekman layer currents arises more from their horizontal spatial extent and variation than from their magnitude alone. Although small in magnitude and in vertical extent (typically the layer extends from the surface to about 100 m depth), the mass transport in the Ekman layer (integrated across the width of the ocean) can be substantial, comparable to the transport of major ocean currents such as the Gulf Stream or the Kuroshio. Spatial variation in the Ekman transport is caused by spatial variation in the wind (the wind stress curl) and results in an

222

exchange of fluid with the ocean interior (Ekman pumping); this exchange of mass induces motion in the ocean interior in order to conserve angular momentum. In regions where the Ekman transport converges, conservation of mass requires that fluid be pumped from the surface Ekman layer into the ocean interior; in regions where the Ekman transport diverges, fluid is pumped into the Ekman layer from below. It is through the vertical velocity or pumping thus generated at the base of the Ekman layer that the wind ultimately forces the ocean interior circulation. Ekman pumping also transports material (nutrients, heat, etc.) from the upper thermocline toward the photic zone and sea surface. The pattern of converging and diverging Ekman layer currents is thus imprinted very strongly upon the patterns of biological productivity as well as upon the strength and shape of the major oceanic current gyres.

Ekman’s Theory of the Wind-driven Currents In 1905 Ekman published a simple but elegant theory to explain F. Nansen’s observations of ice movements. Nansen was an oceanographer and Arctic explorer; he observed that wind blowing over ice floes caused them to drift at an angle of 20–401 to the right of the wind rather than downwind, and he correctly surmised that the earth’s rotation was causing the ice to move at an angle to the wind. Ekman was the first to derive the equations that describe these surface-layer currents, and he used the solution to explain Nansen’s observation. Ekman assumed that the direct influence of the wind was confined to a surface layer, a frictional boundary layer approximately 10–100 m deep, where a steady wind stress was balanced by the Coriolis force. The Coriolis force is an apparent force due to the earth’s rotation; it cannot set the fluid in motion, but it can act to change the motion over timescales of days or longer (i.e., timescales on the order of the earth’s rotation period). Ekman’s key assumptions were: (1) a steady wind blowing over the ocean surface, far from any coast, (2) a flat ocean surface, (3) a constant water density, and (4) a frictional force acting in the surface boundary layer that matched the wind stress at the sea surface and decayed to zero at the bottom of the surface layer. (An inviscid assumption, i.e., neglect of frictional forces, is valid for most of the ocean except near


EKMAN TRANSPORT AND PUMPING

boundaries.) By making these assumptions Ekman arrived at a greatly simplified theoretical ocean, yet he retained the physics required to explain the winddriven surface currents. For example, if the sea surface is horizontal and the density is uniform, the pressure at any depth is constant, and there will be no pressure-driven flows. In the real ocean, the tilt of the sea surface and the internal horizontal density gradients result in flows due to the pressure-gradient force, and separating these pressure-driven flows from wind-driven currents is the main challenge in direct testing of Ekman’s theory from observations. The Coriolis force is given by the vector crossproduct -

Coriolis force ¼ rf kˆ u;

f ¼ 2 O sin f;

½1

where r is the density of seawater, kˆ is the unit vector in the direction of the local vertical, and u ¼ ðu; vÞ is the vector of east ðuÞ and north ðvÞ horizontal currents. The Coriolis parameter f is twice the magnitude of the vertical component of the Earth’s rotation vector O (2p radians per sidereal day) and f is the latitude. Ekman’s steady momentum balance is between the Coriolis force and the vertical gradient of the frictional stress t ¼ ðtx ; ty Þ: -

rf kˆ u ¼

-

@t @z

½2

At the sea surface the frictional stress equals the wind stress t0 . Bulk formulae parameterize the wind stress in terms of a wind velocity at 10 m above the sea surface U10 , the air density rair , and a drag coefficient CD : t0

- - ¼ rair CD U10 U10

½3

A wind speed of 10 m s1 corresponds to a stress of about 0.1 N m2. Ekman’s assumption that the frictional force acted only in the surface layer allowed him to estimate the volume transport within the layer by integrating from the sea surface, where the wind stress was known, to the unknown depth H where the frictional stress vanished by assumption. The Ekman transport relation is given by Mx ¼

ty0 ; rf

My ¼

tx0 rf

½4

where Mx is the eastward component of Ekman transport and My is the northward component. One of the surprising results of the theory is that the net transport is at right angles to the wind direction and depends only on the wind stress at the sea surface.

223

Most notably, the transport does not depend on the details of how the stress gets transferred through the y Ekman layer. The northward wind component t0 forces the eastward Ekman transport Mx ; and the eastward wind component tx0 forces the northward Ekman transport My : The sign (to the right/left) of the wind depends on the sign of the Coriolis parameter f, which is positive/negative in the Northern/ Southern Hemisphere, respectively. The transport is in units of m2 s1, and it can be interpreted as the rate at which the volume (per unit width) of water in the Ekman layer is moving. The qualifier ‘per unit width’ emphasizes that this is the transport at a single point location; in practice, one is usually interested in integrating the Ekman transport over some distance, such as a latitude or longitude band, to see the total volume of water in the Ekman layer that is transported across that latitude or longitude. Spatial variation in the wind and therefore in the Ekman transport results in local convergences and divergences within the surface layer. The vertical velocity (Ekman pumping) wH that results from convergence or divergence in the Ekman layer is derived by integrating the mass conservation equation over the Ekman layer: y @ t0 =rf @ tx0 =rf wH ¼ @x @y

½5

Thus the Ekman pumping is given by the spatial derivative or curl of the wind stress (divided by rf ); it depends on the spatial variation of the wind rather than its magnitude. Ekman’s theory also predicted the currents within the surface layer. His solution for the velocity structure, the Ekman spiral, is the least robust of his results since it depends critically on how one assumes the stress that the wind exerts on the surface is transferred downward by friction and mixing. Ekman was the first to acknowledge that the wind momentum is transferred to ocean currents through turbulent mixing. He modeled it as a diffusion process, exactly analogous to molecular diffusion, but with an effective kinematic viscosity (eddy viscosity) many orders of magnitude larger than molecular viscosity. Turbulent eddies are much more efficient than molecular diffusion at stirring the fluid and hence mixing the wind momentum. For example, a wind of 10 m s1 would require more than a day to penetrate the top meter of the surface layer if molecular diffusion were the sole process acting to mix the wind momentum. In fact, the wind momentum is observed to mix down by tens of meters within hours. Although turbulence is clearly the dominant process in creating the wind-mixed layer, the detailed



_

16 0

12 8

_2 4 _4

Mean wind _ (m s 1) 0

_6 _8

_6

_4 _2 0 _ East velocity (cm s 1)

2

4

4

½6 2

y

tx t0 V ¼ p0ffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ð2Av jf jÞ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi The Ekman depth, DE ¼ ð2Av =jf jÞ, is the depth over which the amplitude decays by 1/e and over which the velocity vector rotates by one radian. Note that the predicted surface current lies at an angle of 451 to the wind direction. Ekman spirals have been observed in the laboratory, where the appropriate viscosity is the molecular value n. More limited observations are available in the open ocean, because of the difficulty in acquiring observations with the requisite vertical resolution and because of the difficulty in separating the wind-driven flow from pressure-driven flow. Ocean spirals have been observed (Figure 1); they tend to be flatter than Ekman’s theory predicts, due to the effect of other processes such as stratification and the diurnal cycling of the mixed layer depth. Regardless of these details, integrating the Ekman spiral (eqn [6]) over the surface layer yields the more general Ekman transport result (eqn [4]).

Ekman Transport and Pumping One of the remarkable results of Ekman’s theory is the Ekman transport relation (eqn [4]), which allows the surface layer transports to be predicted from the wind field. Ship observations of wind have been made over a much longer period of time and over a much greater area of the ocean than have direct measurements of ocean currents, and satellites presently provide global coverage of the wind field. Wind velocity is the quantity that is typically measured, at heights from 10 to 30 m above the sea surface. Wind stress is proportional to the wind speed squared and in the same direction as the wind velocity.

_

y

tx0 þ t0 Vþ ¼ pffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi r ð2Av jf jÞ;

2

(A)

u ¼ expðz=DE Þ½Vþ cosðz=DE Þ þ V sinðz=DE Þ v ¼ expðz=DE Þ½Vþ sinðz=DE Þ V cosðz=DE Þ

4

North velocity (cm s 1)

structure of the turbulent boundary layer remains an active research question today (see Upper Ocean Mixing Processes). To solve for the currents, Ekman assumed a constant eddy viscosity Av in place of the molecular viscosity n. The value of n for sea water is approximately 106 m2 s1 and applies for smooth laminar flow. The magnitude and structure of a turbulent eddy viscosity An are not well known, but scaling arguments suggest that its magnitude may be as large as 101 m2 s1. The current structure is a spiral (Figure 1) that decays in amplitude and rotates clockwise with increasing depth (z negative):

North velocity (cm s 1)

224

0

12

16

8 _2

4 0

_4

Mean wind _ (m s 1)

_6 _8

_6

(B)

_4 _2 0 _ East velocity (cm s 1)

2

4

Figure 1 Comparison of an observed wind-forced velocity spiral and the theoretical Ekman spiral calculated for the same wind in the Northern Hemisphere. Each current vector is the timeaveraged velocity at a particular depth; the first five depths are labelled. Depth units are m, current velocity units are cm s1, and wind velocity units are m s1. (A) The theoretical Ekman spiral: the surface current lies 451 to the right of the wind, and the rate of amplitude decay and the rate of rotation with depth are both set by the constant eddy viscosity (0.014 m2 s1), chosen to match the rate of amplitude decay of the observed spiral. (B) Observed velocity spiral from averaged current observations from a location about 400 km offshore of the US west coast from the Eastern Boundary Currents Experiment. The currents spiral to the right of the wind direction and decay with depth, as predicted by the theory. However, the surface current lies more than the predicted 451 to the right of the wind, and the current amplitude decreases at a faster rate than it turns to the right, resulting in a flatter spiral than the theory predicts. These differences are due to unmodeled processes.

A first step in calculating wind stress from velocity measurements is to determine the wind speed at 10 m height above the sea surface, using a model of the wind profile versus height. The stress at the sea surface can then be calculated using eqn [3] that parameterizes the stress using the 10-m wind speed and a drag coefficient that depends on the speed and other



properties of the air–sea interface such as air and sea temperature and relative humidity. Direct estimation of the wind stress requires measuring the turbulent eddies in the atmospheric boundary layer above the sea surface. These measurements are more difficult to make and hence are not made routinely; however, they are critical to improving and validating the bulk formulae. The Ekman transport relation predicts a transport at right angles to the wind stress, in direct proportion to the wind and inversely proportional to the Coriolis parameter. The theory cannot be applied at the equator where f ¼ 0. For a wind stress of 0.1 N m2, and a seawater density of about 1025 kg m3, the transport per unit width at a latitude of 301 is 1.3 m2 s1 and at a latitude of 101 it is 3.4 m2 s1. Integrated across an ocean basin, the Ekman transport can be quite large. The Atlantic Ocean at 101N is about 4000 km across. The mean wind stress is of order 0.1 N m2, and the predicted Ekman transport is 15 106 m3 s1. This transport is about one-half the transport measured for the Gulf Stream where it passes through the Straits of Florida. Across the same latitude in the Pacific, the Ekman transport is estimated to be about 50 106 m3 s1, not because the winds are stronger but because the Pacific is

90˚N

225

about three times the width of the Atlantic at that latitude. The wind stress is communicated to the ocean interior through Ekman pumping, and therefore the spatial variation of the Ekman transport is just as important as its magnitude. The large-scale pattern of the wind is one of alternating bands of easterlies and westerlies (Figure 2A). The equator is spanned by a broad band of easterly winds known as the trade winds. The magnitude of these winds decreases with latitude, reaching a minimum about 301 or so away from the equator. Still further north/south lies the band known as the prevailing westerlies, with a maximum at about 501 latitude. The polar easterlies are the most poleward band of winds. This overall pattern dominates in both hemispheres of the Pacific and Atlantic Oceans. The Ekman transport relation implies a corresponding pattern of convergences and divergences in the Ekman transport that should be detectable through the Ekman pumping of ocean current gyres (Figure 2C). For example, the theory suggests that the persistent easterly trade winds in the tropics will force a poleward Ekman transport; the midlatitude westerlies will force an equatorward Ekman transport. The resulting Ekman transport convergence at subtropical latitudes should result in

90˚N

90˚N

Polar easterlies

Subpolar gyre

+

(sea surface low)

(upwelling) Prevailing westerlies

Subtropical gyre

_

(sea surface high)

(downwelling) Trade winds 0˚

0˚

Equator (A)

(B)

(C)

Figure 2 Schematic of the relation between zonal wind stress, Ekman transport and pumping, and ocean current gyres. (A) Zonal wind stress for the Northern Hemisphere, with easterly wind stress near the equator and the pole, and westerly wind stress at midlatitudes. (B) The meridional Ekman transport that results from the zonal wind stress. The magnitude is proportional to the wind stress, and the direction is orthogonal to the wind direction. The spatial variation in the wind stress results in Ekman convergences and divergences and Ekman pumping. (C) The wind-driven oceanic current gyres. Ekman convergence in the subtropics results in a seasurface high and a clockwise-current gyre with downwelling at the gyre center. The Ekman divergence at subpolar latitudes results in a sea-surface low and an anticlockwise-current gyre with upwelling at the gyre center.


226


a thick warm layer of water above the main thermocline, an Ekman pumping that supplies fluid to the ocean interior, and, through conservation of angular momentum, a clockwise/anticlockwise general circulation in the northern/southern hemisphere. Such gyres are major observed features of the subtropical oceans. Their existence is an indirect confirmation of Ekman’s theory. It is through the pumping thus generated at the base of the Ekman layer that the wind ultimately forces the ocean interior. If the wind were spatially uniform, the winddriven currents and mass transport would remain largely confined within the shallow surface Ekman layer and would have a negligible role in driving ocean circulation. Note that using the wind to estimate the Ekman transport via eqn [4] gives no indication of how deep it extends. Direct verification, from measurements of both wind and currents, is required in order to determine the details of the turbulent mixing of the wind momentum, such as the maximum depth of frictional influence of the wind, and the importance of other processes such as time dependence and stratification. This direct confirmation eluded oceanographers until quite recently, because of the difficulty in making accurate measurements in the harsh surface zone and in separating the wind-driven flow from other pressure-driven flows. Also, since the wind is not steady, appropriate timescales for averaging need to be determined. Direct verification of the integrated Ekman transport has been made from measurements at tropical latitudes 8–111N in the Atlantic, Pacific, and Indian Oceans, where the Ekman transport is large. High vertical resolution moored measurements at midlatitudes in the Atlantic and the Pacific have verified the Ekman transport relation for moderate winds, with transport per unit width of about 1 m2 s1. Ekman transport and pumping have far-reaching implications, from the mechanism of momentum transfer from wind to water to the maintenance of the large-scale wind-driven oceanic current gyres. Although the original theory was developed for a steady wind, the phenomena of Ekman transport and pumping occur on much shorter timescales, e.g., the scale of synoptic storms. However, Ekman transport and pumping are most effective in driving large-scale ocean circulation when the wind is steady over long periods of time, because then the effects can accumulate.

ind

W

Ekman transport

Ekman layer

Upwelling Onshore flow

Figure 3 Diagram of upwelling (Ekman divergence) at a coast. The wind blows parallel to the coast (in a direction out of the diagram in the Northern Hemisphere), forcing a net Ekman transport offshore. This offshore transport is compensated by onshore flow at depth and upwelling.

uniform wind blowing parallel to a coast can also cause Ekman pumping. In the Northern Hemisphere, Ekman divergence occurs when the wind blows parallel to a coastline on its left. For example, during spring and summer the mean winds along the west coast of North America are southward. Associated with these winds is a net westward Ekman transport. This offshore transport of mass in the surface layer is balanced by an onshore flow at depth and an Ekman pumping of water from depth into the surface layer (Figure 3). This phenomenon is known as coastal upwelling, and it occurs seasonally along the west coasts of continents in both hemispheres. Prolonged upwelling can provide a continuous source of cold, nutrient-rich water to the surface layer euphotic zone, replacing the nutrient-depleted water that is transported offshore. Many of the ocean’s important fisheries are concentrated in upwelling zones. The reverse phenomenon, convergence from onshore Ekman transport and downward pumping at the coast also occurs and is called downwelling.

Summary Coastal Upwelling Although spatial variation in the wind field is the cause of Ekman pumping in the open ocean, a spatially

Ekman’s theory of wind-driven currents, now almost 100 years old, is a cornerstone of physical oceanography. The two essential features of the theory – that



the dominant momentum balance for steady winddriven currents is between the wind stress and Coriolis acceleration and that wind stress is transmitted through the upper ocean as a turbulent momentum flux – are now well established by observation. The most important predictions of the theory, the transport relation and Ekman pumping, form the essential link between the winds and our understanding of the winddriven ocean circulation and is also established by observation. Yet there are other aspects of Ekman’s theory, most notably the eddy diffusivity, that are still unsettled in the sense that they cannot be derived strictly from a more general theory, and neither is there strong support for eddy diffusivity from observation. This may seem to indicate unusually slow progress on this important point, and yet the uncertainty surrounding eddy diffusivity is no more than a reflection of the tortuous development of turbulent transfer theory generally. It seems likely that succeeding editions of this encyclopedia will relate a similar story on this point, at least.

227

Further Reading Chereskin TK (1995) Direct evidence for an Ekman balance in the California Current. Journal of Geophysical Research 100: 18261--18269. Ekman VW (1905) On the influence of the earth’s rotation on ocean-currents. Arkiv fo¨r Matematik, Astronomi och Fysik 2: 1--52. Gill AE (1982) Atmosphere–Ocean Dynamics. New York: Academic Press. Pond S and Pickard GL (1983) Introductory Dynamical Oceanography, 2nd edn. Oxford: Pergamon. Price JF, Weller RA, and Schudlich RR (1987) Wind-driven currents and Ekman transport. Science 238: 1534--1538.

See also Internal Tidal Mixing. Meddies and Sub-Surface Eddies. Mesoscale Eddies. Seabird Population Dynamics. Upper Ocean Mixing Processes. Wave Generation by Wind. Wind- and Buoyancy-Forced Upper Ocean. Wind Driven Circulation.


˜ O SOUTHERN OSCILLATION (ENSO) EL NIN K. E. Trenberth, National Center for Atmospheric Research, Boulder, CO, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 815–827, & 2001, Elsevier Ltd.

Introduction A major El Ninõ began in April of 1997 and continued until May 1998. It has been labeled by some as the ‘El Ninõ of the century’ as it was certainly the biggest on record by several measures. It brought with it many weather extremes and unusual weather patterns around the world. Moreover, the event was predicted by climate scientists and received unprecedented news coverage, so that the term ‘El Ninõ’ became part of the popular vernacular. Many things were blamed on El Ninõ, and some of them indeed were influenced by El Ninõ, although in some instances, the connection was, at best, tenuous. Although El Ninõ may be relatively new to the public, it has been known to scientists, at least in some respects, for decades and even centuries. El Ninõ refers to the exceptionally warm sea temperatures in the tropical Pacific, but it is linked to major changes in the atmosphere through the phenomenon known as the Southern Oscillation (SO), in particular, so that the whole phenomenon is called El Ninõ–Southern Oscillation (ENSO) by scientists. This article outlines the current understanding of ENSO and the physical connections between the tropical Pacific and the rest of the world. ENSO Events

El Ninõs are not uncommon. Every three to seven years or so, a pronounced warming occurs of the surface waters of the tropical Pacific Ocean. The warmings take place from the International Dateline to the west coast of South America and result in changes in the local and regional ecology, and are clearly linked with anomalous global climate patterns. In 1997, the warming can be seen by comparing the sea surface temperatures (SSTs) in December 1997 at the peak of the 1997/98 event with those a year earlier (Figure 1). As well as the total SST fields, this figure also displays the departures from average. The warmings have come to be known as ‘El Ninõ events’. Historically, ‘El Ninõ’ referred to

228

the appearance of unusually warm water off the coast of Peru, where it was readily observed as an enhancement of the normal warming about Christmas (hence Ninõ, Spanish for ‘the boy Christchild’) and only more recently has the term come to be regarded as synonymous with the basinwide phenomenon. The oceanic and atmospheric conditions in the tropical Pacific are seldom close to average, but instead fluctuate somewhat irregularly between the warm El Ninõ phase of ENSO, and the cold phase of ENSO consisting of basinwide cooling of the tropical Pacific, dubbed ‘La Ninã events’ (‘La Ninã’ is ‘the girl’ in Spanish). The most intense phase of each event typically lasts about a year. The SO is principally a global-scale seesaw in atmospheric sea level pressure involving exchanges of air between eastern and western hemispheres (Figure 2) centered in tropical and subtropical latitudes with centers of action located over Indonesia and the tropical South Pacific Ocean (near Tahiti). Thus the nature of the SO can be seen from the inverse variations in pressure anomalies (departures from average) at Darwin (12.41S 130.91E) in northern Australia and Tahiti (17.51S 149.61W) in the South Pacific Ocean (Figure 3) whose annual mean pressures are strongly and significantly oppositely correlated. Consequently, the difference in pressure anomalies, Tahiti–Darwin, is often used as a Southern Oscillation Index (SOI). The sequences of swings in the SOI shown in Figure 3 are discussed below in conjunction with those of SST. Higher than normal pressures are characteristic of more settled and fine weather, with less rainfall, whereas lower than normal pressures are identified with ‘bad’ weather, more storminess and rainfall. So it is with the SO. Thus for El Ninõ conditions, higher than normal pressures over Australia, Indonesia, southeast Asia, and the Philippines signal drier conditions or even droughts. Dry conditions also prevail at Hawaii, parts of Africa, and extend to the northeast part of Brazil and Colombia. On the other end of the seesaw, excessive rains prevail over the central and eastern Pacific, along the west coast of South America, parts of South America near Uruguay, and southern parts of the United States in winter (cf. Figure 2) often leading to flooding. When the pressure pattern in Figure 2 reverses in sign, as for La Ninã, the regions favored for drought in El Ninõ tend to become excessively wet, and vice versa.



_3

_2 _1 0

1 2

3

4

0

(D)

90S

30E

0

30E

Sea surface temperature 1 degree (smoothed)

_3

_2

_1

0

1

2

3

4

0

(˚C)

60E 90E 120E 150E 180 150W 120W 90W 60W 30W Contour from _ 3 to 4 by 1

Anomalous sea surface temperature 1 degree (smoothed)

2 4 6 8 1012 14 16 18 2022 24262830

0

(˚C)

60E 90E 120E 150E 180 150W 120W 90W 60W 30W Contour from 2 to 30 by 2

December 1997

0

December 1997

Figure 1 Monthly mean sea surface temperatures in 1C for December 1996 (A, C) and 1997 (B, D), before and during the peak of the 1997–98 El Ninõ event. (A) and (B) show the actual SSTs and (C) and (D) show the anomaly, defined as the departure from the mean for 1950–79 with contour interval 21C (top) and 11C (bottom). Based on data from US National Oceanic and Atmospheric Administration.

(C)

60E 90E 120E 150E 180 150W 120W 90W 60W 30W Contour from _ 3 to 4 by 1

60S

60S

30E

30S

30S

0

0

0

90S

30N

30N

120E

60N

(˚C)

60N

Anomalous sea surface temperature 1 degree (smoothed)

(B)

90N

December 1996

2 4 6 8 1012 14 16 18 2022 24262830

90N

(A)

Contour from 2 to 30 by 2

90S

60S

60S

0

30S

30S

30E 60E 90E 120E 150E 180 150W 120W 90W 60W 30W

0

0

0

30N

30N

90S

60N

60N

(˚C) 90N

Sea surface temperature 1 degree (smoothed)

90N

December 1996

˜ O SOUTHERN OSCILLATION (ENSO) EL NIN 229

˜ O SOUTHERN OSCILLATION (ENSO) EL NIN

230

NCEP reanalysis SLP and Darwin SLP Annual (May to April) mean 1958 _ 1998

Correlation

(× 10)

90N 0

0

3

0

60N

0

3

3

0

0

0 _3

6

6

3

3

_3

3

30N

0

_3

3

0

_6

9

3

6 0

3

3

3

30S 3 0 _6

60S

_3 _3

_3 3

0

_3

90S 0

30E

60E

90E

120E

180 150W 150E Contour from _ 7.5 to 9 by 1.5

120W

90W

60W

30W

0

Figure 2 Map of correlation coefficients ( 10) of the annual mean sea level pressures with that at Darwin. It shows the Southern Oscillation pattern in the phase corresponding to El Ninõ events; during La Ninã events the signs are reversed. Values exceeding 0.6 in magnitude are stippled or hatched.

The Tropical Pacific Ocean–Atmosphere System The distinctive pattern of average sea surface temperatures in the Pacific Ocean sets the stage for ENSO events. The pattern in December 1996 (Figure 1) is sufficiently close to average to illustrate the main points. One key feature is the ‘warm pool’ in the tropical western Pacific, where the warmest ocean waters in the world reside and extend to depths of over 150 m with values at the surface 4281C. Other key features include warm waters north of the equator from about 5 to 151N, much colder waters in the eastern Pacific, and a cold tongue along the equator that is most pronounced about October and weakest in March. The warm pool migrates with the sun back and forth across the equator but the distinctive patterns of SST are brought about mainly by the winds (Figure 4). The existence of the ENSO phenomenon is dependent on the east–west variations in SSTs (Figure 1) in the tropical Pacific, and the close links with sea-level pressures (Figure 3) and thus surface winds in the tropics (Figure 4), which in turn determine the major areas of rainfall (Figure 5). The temperature of the surface waters is readily conveyed to the overlying atmosphere and because warm air is less dense

it tends to rise whereas cooler air sinks. As air rises into regions where the air is thinner, the air expands, causing cooling and therefore condensing moisture in the air, which produces rain. Low sea-level pressures are set up over the warmer waters while higher pressures occur over the cooler regions in the tropics and subtropics, and the moisture-laden winds tend to blow toward low pressure so that the air converges, resulting in organized patterns of heavy rainfall. The rain comes from convective cloud systems, often as thunderstorms, and perhaps as tropical storms or even hurricanes, which preferentially occur in the ‘convergence zones’. Because the wind is often light or calm right in these zones, they have previously been referred to as the ‘doldrums’. Of particular note are the Intertropical Convergence Zone (ITCZ) and the South Pacific Convergence Zone (SPCZ) (Figure 5) which are separated by the equatorial dry zone. These atmospheric climatological features play a key role in ENSO as they change in character and move when SSTs change. The rainfall patterns in the tropics can be illustrated by quantities sensed from satellite (Figure 5). There is a strong coincidence between the patterns of SSTs and tropical convection throughout the year, although there is interference from effects of nearby land and monsoonal circulations. The strongest



231

2.5 2.0 1.5

Anomaly (mb)

1.0 0.5 0.0 _ 0.5 _ 1.0 _ 1.5 _ 2.0 (A)

_ 2.5 5 4

Normalized anomalies

3 2 1 0 _1 _2 _3 _4 _5 1950

1955

1960

1965

1970

(B)

1975 Year

1980

1985

1990

1995

2000

Figure 3 (A) Time series of anomalies in sea level pressures at Darwin (solid line) and Tahiti (light line) from 1950 to 1998 smoothed to remove fluctuations of period less than about a year. (B) Corresponding time series of the Southern Oscillation Index, normalized Tahiti minus Darwin pressures, in standardized units.

seasonal migration of rainfall occurs over the tropical continents, Africa, South America and the Australian–Southeast Asian–Indonesian maritime region. Over most of the Pacific and Atlantic the ITCZ remains in the Northern Hemisphere year round, with convergence of the tradewinds favored by the presence of warmer water. In the subtropical Pacific, the SPCZ also lies over waters warmer than about 271C. The ITCZ is weakest in January in the Northern Hemisphere when the SPCZ is strongest in the Southern Hemisphere. The surface winds (Figure 4) drive surface ocean currents which determine where the surface waters flow and diverge, and thus where cooler nutrient-rich waters upwell from below. Because of the Earth’s rotation, easterly winds along the equator deflect

currents to the right in the Northern Hemisphere and to the left in the Southern Hemisphere and thus away from the equator, creating upwelling along the equator. Thus the winds largely determine the SST distribution, the differential sea levels and the heat content of the upper ocean. The presence of nutrients and sunlight in the cool surface waters along the equator and western coasts of the Americas favors development of tiny plant species (phytoplankton), which in turn are grazed on by microscopic sea animals (zooplankton) which provide food for fish. Temperatures in the upper ocean are measured by about 70 instrumented buoys moored to the bottom of the ocean throughout the tropical Pacific (Figure 6). Thus for December 1996 and 1997 the


232


December 1997

T42

m/s

90N

60N 30N 0 30S 60S 90S 0 (A)

30E 60E 90E 120E 150E 180 150W 120W 90W 60W 30W Contour from 5 to 5 by 5

December 1996

T42

0

m/s

90N

60N 30N 0 30S 60S 90S 0 (B)

30E 60E 90E 120E 150E 180 150W 120W 90W 60W 30W 0 Contour from 5 to 5 by 5 _ m/s Dec 1997 Dec 1996 T42

90N

60N 30N 0 30S 60S 90S 0 (C)

30E

60E 90E 120E 150E 180 150W 120W 90W 60W 30W Contour from 3 to 9 by 3 8

0

Reference Vector Figure 4 Mean surface winds for December 1996 (A) and 1997 (B) from global analyses by the European Centre for Medium Range Forecasts, and their difference (C). Shading indicates 5 m s 1 for (A) and (B) and 3 m s 1 for (C). The reference vector is plotted at bottom.



234

30˚N

TAO Array 15˚N 0˚ 15˚S Atlas Current meter

30˚S 120˚E

140˚E

160˚E

180˚

160˚W 140˚W 120˚W 100˚W

80˚W

Figure 6 A center piece of the Pacific El Ninõ observing system is an array of buoys in the tropical Pacific moored to the ocean bottom known as the TAO (Tropical Atmosphere–Ocean) array. The latter is maintained by a multinational group spearheaded in the United States by the National Oceanic and Atmospheric Administration’s Pacific Marine Environmental Laboratory (PMEL). Each buoy has a series of temperature measurements on a sensor cable on the upper 500 m of the mooring, and on the buoy itself are sensors for surface wind, sea surface temperature, surface air temperature, humidity, and a transmitter to satellite. Observations are continually made, averaged into hourly values, and transmitted via satellite to centers around the world for prompt processing. Right, an ATLAS (Autonomous Temperature Line Acquisition System) buoy. Courtesy PMEL.

140˚E

180

140˚W

0

140˚E

100˚W

0

180

16

300

12 8 4

10

1

4 0

0

_4

400

_8

500

_ 12 Anomalies

12

0

8

0

0

0 Means

100

_3

_

2

5

3

5

6

10

_ 1

8 4

0

300

8 4

10

4

1

2

12

500

12

200

20 16

300

(B)

100

14 12

400

Means 0

24

200

0

500 (A)

Depth (m)

Depth (m)

20

400

(C)

100

24 14 12

28

28

Depth (m)

Depth (m)

28

200

100˚W 32

28 100

140˚W

32

200 0 300 400

500 (D) Anomalies

_4 0

_8 _ 12

Figure 7 Zonal cross sections along the equator across the Pacific Ocean (21N to 21S) from the TAO array temperatures for December 1996 (A, C) and 1997 (B, D). (A, B) total field; (C, D) anomalies. Contours every 1C.

temperature structure throughout the equatorial region can be mapped (Figure 7). The heat content of the upper ocean depends on the configuration of the thermocline (the region of sharp temperature gradient within the ocean separating the well-mixed surface layers from the cooler abyssal ocean waters). Normally the thermocline is deep in the western tropical Pacific (150 m) and sea level is high as waters driven by the easterly tradewinds pile up. In the

eastern Pacific on the equator, the thermocline is very shallow (50 m) and sea level is relatively low. The Pacific sea surface slopes up by about 60 cm from east to west along the equator. The temperatures in December 1996 depict conditions somewhat similar to average, but with signs of the incipient El Ninõ developing in the western tropical Pacific at 100– 150 m depth as a warm anomaly (Figure 7) but no such signs at the surface (Figure 1).



The tropical Pacific, therefore, is a region where the atmospheric winds are largely responsible for the tropical SST distribution which, in turn, is very much involved in determining the precipitation distribution and the tropical atmospheric circulation. This sets the stage for ENSO to occur.

Mechanisms of ENSO Most of the interannual variability in the atmosphere in the tropics and a substantial part of the variability over the extratropics is related and tied together through ENSO. ENSO is a natural phenomenon arising from coupled interactions between the atmosphere and the ocean in the tropical Pacific Ocean, and there is good evidence from cores of coral and glacial ice in the Andes that it has been going on for millennia. During El Ninõ, the tradewinds weaken (Figure 4) which causes the thermocline to become shallower in the west and deeper in the eastern tropical Pacific (Figure 7), while sea level falls in the west and rises in the east by as much as 25 cm as warm waters surge eastward along the equator. Equatorial upwelling decreases or ceases and so the cold tongue weakens or disappears (e.g., as in December 1997, Figure 1) and the nutrients for the food chain are substantially reduced. The resulting increase in sea temperatures (e.g., Figure 1) warms and moistens the overlying air so that convection breaks out, and the convergence zones and associated rainfall move to a new location with a resulting change in the atmospheric circulation (Figure 5). A further weakening of the surface trade winds completes the positive feedback cycle leading to an El Ninõ event. The shifts in the location of the organized rainfall in the tropics and the latent heat released alters the heating patterns of the atmosphere. Somewhat like a rock in a stream of water, the anomalous heating sets up waves in the atmosphere that extend into midlatitudes altering the winds and changing the jet stream and storm tracks (e.g., Figure 8). Note especially the strong westerly jets in the Pacific of both hemispheres, and in the Northern (winter) Hemisphere the jet stream in December 1997 extends into California and across the southern United States, carrying with it disturbances that result in extensive rains. Weaker westerlies exist farther north and so the overall storm track shifts towards the equator in the Pacific. Although the El Ninõs and La Ninãs are often referred to as ‘events’ which last a year or so, ENSO is oscillatory in nature. The ocean is a source of moisture and its enormous heat capacity acts as the flywheel that drives the system through its memory

235

of the past, resulting in an essentially self-sustained sequence in which the ocean is never in equilibrium with the atmosphere. The amount of warm water in the tropics builds up prior to and is then depleted during El Ninõ. During the cold phase with relatively clear skies, solar radiation heats up the tropical Pacific Ocean, the heat is redistributed by currents, with most of it being stored in the deep warm pool in the west or off the equator such as at about 10 or 201N. During El Ninõ, heat is transported out of the tropics within the ocean toward higher latitudes in response to the changing currents, and increased heat is released into the atmosphere mainly in the form of increased evaporation, thereby cooling the ocean. Added rainfall contributes to a general warming of the global atmosphere that peaks a few months after a strong El Ninõ event. It has therefore been suggested that the time scale of ENSO is determined by the time required for an accumulation of warm water in the tropics to essentially recharge the system, plus the time for the El Ninõ itself to evolve. Thus a major part of the onset and evolution of the events is determined by the history of what has occurred one to two years previously. This also means that the future evolution is predictable for several seasons in advance.

Interannual Variations in Climate The subsurface temperature anomalies which eventually developed into the 1997–98 El Ninõ were traceable at least from about September 1996 on the equator in the far western Pacific. However, positive subsurface temperature anomalies in the upper 100– 200 m in the far western Pacific exceeded 11C for all the months of 1996, and so this was not a sufficient predictor. By December 1996 (Figure 7) subsurface temperature anomalies in the vicinity of the equator had grown to exceed 2.51C at 150 m depth and the warm anomaly extended from at least 1401E (the westernmost buoy) to 1401W. However, conditions were still below normal in the eastern Pacific. By December 1997, the subsurface warm anomaly had progressed eastward and amplified to produce positive anomalies exceeding 111C (Figure 7) at about 100 m depth, accompanying the surface SST anomalies exceeding 51C (Figure 1). Also note, however, in December 1997 the subsequent cold anomaly of over 51C in the western equatorial Pacific near 150 m depth, as the warm pool was displaced into the central Pacific. The warm pool in the east continued to diminish with time as the cold anomaly intensified and subsequently spread all the way across the Pacific as part of the signature of the La Ninã that


236


_

( × 106 m2 s 1)

December 1996 90N 60N

_ 120

_ 120

_ 120

_ 120

_ 80

_ 80

_ 80

_ 80

30N 0

0 0

0

0

0 40

30S

40

60S

120

90S (A) 0

30E

90E

60E

40

40 120

120

120E 150E

150W 120W 90W

180

60W

0

30W

_1

( × 10 m s ) 6

December 1997

2

90N _120 _ 80

_ 120

60N

_ 80 30N

_ 40

0

_ 120 0

0

0

0

120

0 40

30S

40

80 80 60S

120

80

80

120

120

120

120

90S (B) 0

40

40

30E

60E

90E

120E 150E

180

150W 120W

90W

60W

0

30W

_1

( × 106 m2 s )

Dec ’97 _ Dec ’96 90N 0 60N

0

0

0

_ 40

40

0

0 30N 0

_ 40 0

0

0

0

0

40 0

0

0

30S 0

0 0

60S

0

0 0

0 90S (C) 0

0 30E

60E

90E

120E 150E

180

150W 120W 90W

60W

30W

0

Figure 8 Flow patterns (winds) at the jet stream level, 10 km or so above the surface of the Earth, for December 1996 (A) and 1997 (B), and their differences (C) based on global analyses by the European Centre for Medium Range Prediction. The contours are stream-lines at the 200 mb level and the region where winds exceed 40 m s 1 (20 m s 1 for (C)) are stippled to indicate the jet streams. El Ninõ conditions favor a more vigorous jet stream and storm track from the Pacific to the Gulf of Mexico, bringing heavy rains to the southern United States. Note similar effects in the South Pacific near New Zealand.



began about June 1998, although it did not develop to a full fledged event until later in the year. A key aspect of these changes was the obvious loss of heat content throughout the equatorial Pacific beginning late 1997 and continuing through 1998. The evolution of SST in several recent ENSO events after 1950 is shown in Figure 9 for two regions. The region of the Pacific Ocean which is most involved in ENSO is the central equatorial Pacific, especially the area 51N to 51S, 1701E to 1201W, whereas the traditional El Ninõ region is along the coast of South America. The latter is less important for the global changes in weather patterns but is certainly important locally. Variations in both regions are closely related but differ in detail from event to event in relative amplitudes and sequencing. For SSTs, the departure from average required for an El Ninõ is 0.51C over the central Pacific region, which is large enough to cause perceptible effects in Pacific rim countries. Larger El Ninõ events have traceable influences over more extensive regions and even globally. The ENSO events clearly identifiable in Figure 9 since 1950 occurred in 1951, 1953, 1957–58, 1963,

237

1965, 1969, 1972–73, 1976–77, 1982–83, 1986–87, 1990–95 and 1997–98. The 1990–95 event might also be considered three modest events where the conditions in between failed to return to below normal so that they merged together. Worldwide climate anomalies lasting several seasons have been identified with all of these events. The 1997–98 event has the biggest SST departures on record, but for the SOI, the El Ninõ event of 1982–83 still holds the record (Figure 3). Each El Ninõ event has its own character. In the El Ninõ winters of 1992–93, 1994–95, and 1997–98, southern California was battered by storms and experienced flooding and coastal erosion (in part aided by the high sea levels). However, in more modest El Ninõs (e.g., 1986–87 and 1987–88 winters) California was more at risk from droughts. Because of the enhanced activity in the Pacific and the changes in atmospheric circulation throughout the tropics, there is a decrease in the number of tropical storms and hurricanes in the tropical Atlantic during El Ninõ. A good example is 1997, one of the quietest Atlantic hurricane seasons on record, whereas the 1990–95 and 1997–98 El Ninõs terminated before

5

SST anomaly (˚C)

4 3 2 1 0 _1 (A) _ 2 5

SST anomaly (˚C)

4 3 2 1 0 _1 _2 1950 (B)

1955

1960

1965

1970

1975 Year

1980

1985

1990

1995

2000

Figure 9 Time series of sea surface temperature (SST) anomalies from 1950 through 1998 relative to the means of 1950–79 for (A) the traditional El Ninõ area: 01–101S 901W–801W, and (B) the region most involved in ENSO 51N–51S, 1701E–1201W.


238


the 1995 and 1998 hurricane seasons which unleashed the Atlantic storms and placed those seasons among the most active on record. The SO has global impacts, however; the connections to higher latitudes (known as teleconnections) tend to be strongest in the winter of each hemisphere and feature alternating sequences of high and low pressures accompanied by distinctive wave patterns in the jet stream (Figure 8) and storm tracks in mid-latitudes. Although warming is generally associated with El Ninõ events in the Pacific and extends, for instance, into western Canada and Alaska, cool conditions typically prevail over the North and South Pacific Oceans. To a first approximation, reverse anomaly patterns occur during the La Ninã phase of the phenomenon. The prominence of the SO has varied throughout the last century (Figure 10). Very slow long-term (decadal) variations are present; for instance SOI values are mostly below the long-term mean after 1976. This accompanies the generally above normal SSTs in the western Pacific along the equator (Figure 9). The decadal atmospheric and oceanic variations are even more pronounced in the North Pacific and across North America than in the tropics and are also clearly present in the South Pacific, with evidence suggesting that they are at least in part forced from the tropics. Although not yet clear in detail, it is likely that climate change associated with increased greenhouse gases in the atmosphere, which contribute to global warming, are changing ENSO, perhaps by expanding the west Pacific warm pool and making for more frequent and bigger El Ninõ events.

Impacts Changes associated with ENSO produce large variations in weather and climate around the world from year to year and often these have a profound impact on humanity because of droughts, floods, heat waves and other changes which can severely disrupt agriculture, fisheries, the environment, health, energy demand, and air quality, and also change the risks of fire. The normal upwelling of cold nutrient-rich and CO2-rich waters in the tropical Pacific is suppressed during El Ninõ. The presence of nutrients and sunlight fosters development of phytoplankton and zooplankton to the benefit of many fish species. Therefore El Ninõ-induced changes in oceanic conditions can have disastrous consequences for fish and seabirds and thus for the fishing and guano industries, for example, along the South American coast. Other marine creatures may benefit so that unexpected harvests of shrimp or scallops occur in some places. Rainfall over Peru and Ecuador can transform barren desert into lush growth and benefit some crops, but can also be accompanied by swarms of grasshoppers, and increases in the populations of toads and insects. Human health is affected by mosquito-borne diseases such as malaria, dengue, and viral encephalitis, and by water-borne diseases such as cholera. Economic impacts can be large, with losses typically overshadowing gains. One assessment placed losses during the 1997–98 El Ninõ event at over US$34 billion, although it does not account for the widespread human suffering and loss of life. An estimate of the economic loss of production in

2.5

Normalized anomalies

2.0 1.5 1.0 0.5 0.0 _ 0.5 _ 1.0 _ 1.5 _ 2.0 _ 2.5 1880 1890 1900 1910 1920 1930 1940 1950 1960 1970 1980 1990 2000 Year Figure 10 Time series of the Southern Oscillation Index (SOI) based solely on observations at Darwin from 1866 to 1998. Also shown in a curve designed to show the multidecadal fluctuations. The zero corresponds to the mean for the first 100 years 1866 to 1965, highlighting the recent trend for more El Ninõs.



Queensland, Australia due to drought in the prolonged warm ENSO phase from 1990 to 1995 is $1 billion (Australian) per year. The La Ninã in 1998 was linked to the extensive, severe, and persistent 1988 North American spring–summer drought, which brought losses of over US$30 billion, as well as major climate anomalies elsewhere over the globe. ENSO also plays a prominent role in modulating carbon dioxide exchanges with the atmosphere. The decrease in outgassing of CO2 during El Ninõ is enough to reduce the build up of CO2 in the atmosphere by 50%. El Ninõ also influences the incidence of fires, which result in more CO2 emissions, while changing rainfall and temperatures over land through the teleconnections, such that CO2 uptake by the terrestrial biosphere is enhanced. In 1997, the strongest drought set in over Indonesia and led to many fires, set as part of activities of farmers and corporations clearing land for agriculture, raging out of control. With the fires came respiratory problems in adjacent areas 1000 km distant and even a commercial plane crash in the area has been linked to the visibility problems. Subsequently in 1998, El Ninõ-related drought and fires evolved in Brazil, Mexico and Florida. Flooding took place in Peru and Ecuador, as usual with El Ninõ, and also in Chile, and coastal fisheries were disrupted. The strong winter 1997–98 Northern Hemisphere jet stream created wet conditions from California to Florida. Normally these storms veer to the north toward the Gulf of Alaska or enter North America near British Columbia and Washington, where they could subsequently link up with the cold Arctic and Canadian air masses and bring them down into the United States. Instead, the pattern was persistently favorable for relatively mild conditions over the northern states such that temperatures averaged over 101C (181F) above normal in February in the Great Lakes area. Similar changes occurred in the Southern Hemisphere and spread downstream to South America. Globally, it seems that the land temperature for February 1998 was relatively higher than average than for any other month on record. The calendar year 1998 was the warmest year on record – going back 1000 years – beating 1997, in no small part because of the El Ninõ influences in both those years.

239

predictions in the tropics of at least six months have been shown to be practicable. For instance, an El Ninõ was predicted for 1997 in late 1996 and it began in April 1997. However, its full extent was not predicted accurately until about mid-1997, about 6 months in advance. Further improvements in the climate observing system and more realistic and comprehensive models provide prospects for further advances. It is already apparent that reliable prediction of tropical Pacific SST anomalies can lead to useful skill in forecasting rainfall anomalies in parts of the tropics. Although there are certain common aspects to ENSO events in the tropics, the effects at higher latitudes are more variable. One difficulty is the vigor of weather systems in the extratropics which can override relatively modest ENSO influences from the tropics. Nevertheless, systematic changes in the jet stream and storm tracks do tend to occur on average, thereby allowing useful predictions to be made in some regions, although with some inherent uncertainty, so that the predictions are couched in terms of probabilities. Skillful seasonal predictions of temperatures and rainfalls have the potential for huge benefits for society, although because the predictability is somewhat limited, a major challenge is to utilize the uncertain forecast information in the best way possible throughout different sectors of society (e.g., crop production, forestry resources, fisheries, ecosystems, water resources, transportation, energy use). The utility of a forecast may vary considerably according to whether the user is an individual versus a group or country. An individual may find great benefits if the rest of the community ignores the information, but if the whole community adjusts (e.g., by growing a different crop), then supply and market prices will change, and the strategy for the best crop yield may differ substantially from the strategy for the best monetary return. On the other hand, the individual may be more prone to small-scale vagaries in weather that are not predictable. Vulnerability of individuals will also vary according to the diversity of the operation: whether there is irrigation available, whether the farmer has insurance, and whether he or she can work off the farm to help out in times of adversity.

See also ENSO and Seasonal Predictions The main features of ENSO have been captured in models, in particular in simplified models which predict the anomalies in SSTs. Lead times for

Carbon Dioxide (CO2) Cycle. El Ninõ Southern Oscillation (ENSO) Models. Evaporation and Humidity. Fisheries and Climate. Heat Transport and Climate.Ocean Circulation. Pacific Ocean Equatorial Currents.


240


Further Reading Glantz MH, Katz RW, and Nicholls N (eds.) (1991) Teleconnections Linking World-wide Climate Anomalies. Cambridge: Cambridge University Press. Glantz MH (1996) Currents of Change: El Ninõ’s Impact on Climate and Society. Cambridge: Cambridge University Press. National Research Council (1996) Learning to Predict Climate Variations Associated with El Ninõ and the Southern Oscillation: Accomplishments and Legacies of the TOGA Program. Washington: National Academy Press.

Philander SGH (1990) El Ninõ, La Ninã, and the Southern Oscillation. London: Academic Press. Suplee C (1999) El Ninõ/La Ninã National Geographic, March. 72–95. Trenberth KE (1997) Short-term climate variations: recent accomplishments and issues for future progress. Bulletin of the American Meteorological Society 78: 1081--1096. Trenberth KE (1997) The definition of El Ninõ. Bulletin of the American Meteorological Society 78: 2771--2777. Trenberth KE (1999) The extreme weather events of 1997 and 1998. Consequences 5(1): 2--15.


˜ O SOUTHERN OSCILLATION (ENSO) EL NIN MODELS S. G. Philander, Princeton University, Princeton, NJ, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 827–832, & 2001, Elsevier Ltd.

Introduction The signature of El Ninõ is the interannual appearance of unusually warm surface waters in the eastern tropical Pacific Ocean. That area is so vast that the effect on the atmosphere is profound. Rainfall patterns are altered throughout the tropics – some regions experience floods, others droughts – and even weather patterns outside the tropics are affected significantly. From an atmospheric perspective, these various phenomena are attributable to the change in the sea surface temperature pattern of the tropical Pacific. Why does the pattern change? Whereas sea surface temperatures depend mainly on the incident solar radiation over most of the globe, the tropics are different. There the winds are of primary importance because of the shallowness of the thermocline, the thin layer of large temperature gradients, at a depth of approximately 100 m, that separates warm surface waters from the cold water at depth. The winds, by causing variations in the depth of the thermocline, literally bring the deep, cold water to the surface in regions where the thermocline shoals. For example, the trade winds that drive warm surface waters westward along the equator expose cold, deep water to the surface in the eastern equatorial Pacific. A relaxation of the winds, such as occurs during El Ninõ,

permits the warm water to flow back eastward. The changes in the winds are part of the atmospheric response to the altered sea surface temperatures. This circular argument – the winds are both the cause and consequence of sea surface temperature changes – suggests that interactions between the ocean and atmosphere are at the heart of the matter. Those interactions give rise to a broad spectrum of natural modes of oscillation. This result has several important implications. One is that El Ninõ, even though we tend to regard him as an isolated phenomenon that visits sporadically, is part of a continual fluctuation, known as the Southern Oscillation. El Ninõ corresponds to one phase of this oscillation, the phase during which sea surface temperatures in the eastern tropical Pacific are unusually high. La Ninã is the name for the complementary phase, when temperatures are below normal. (Very seldom are temperatures ‘normal’, as is evident in Figure 1.) To ask why El Ninõ, or La Ninã, occur, is equivalent to asking why a pendulum spontaneously swings back and forth. Far more interesting questions concern the factors that determine the period and other properties of the oscillation, and the degree to which it is self-sustaining or damped. Only strongly damped oscillations, that disappear at times, require a trigger to get going again. Hence a search for the disturbance that triggered a particular El Ninõ is based on an implicit assumption, which may not be correct at all times, that the Southern Oscillation was damped and had disappeared for a while. The tools for predicting El Ninõ are coupled models of the ocean and atmosphere; each provides boundary conditions for the other, sea surface temperatures in the

Temperature (°C)

28 27 26 25 24 1900

1910

1920

1930

1940

1950

1960

1970

1980

1990

2000

Year Figure 1 Surface temperature fluctuations at the equator (in 1C) in the eastern equatorial Pacific (after removal of the seasonal cycle) over the past 100 years. Temperature maxima correspond to El Ninõ, minima to La Ninã. The smoothly varying dashed line is a 10-year running mean.


241

242

˜ O SOUTHERN OSCILLATION (ENSO) MODELS EL NIN

one case, the winds in the other case. The following discussion concerns first atmospheric then oceanic models, and finally coupled models.

The Atmosphere The atmospheric circulation in low latitudes corresponds mainly to direct thermal circulations driven by convection over the regions with the highest surface temperatures. Moisture-bearing trade winds converge onto these regions where the air rises in cumulus towers that provide plentiful rainfall locally. The three main convective zones – they can easily be identified in satellite photographs of the Earth’s cloud cover – are over the Congo and Amazon River basins, and the ‘maritime continent’ of the western equatorial Pacific, south-eastern Asia, and northern Australia. The latter region includes an enormous pool of very warm water that extends to the vicinity of the dateline. Much of the air that rises there subsides over the relatively cold eastern equatorial Pacific where rainfall is minimal; the deserts along the coasts of Peru and California in effect extend far westward over the adjacent ocean. The only region of warm surface waters and heavy rainfall in the eastern tropical Pacific is the doldrums, also known as the Intertropical Convergence Zone, a narrow band between 31 and 101N approximately, onto which the south-east and north-east trade winds converge. A warming of the eastern tropical Pacific, such as occurs during El Ninõ, amounts to an eastward expansion of the pool of warm waters over the western Pacific and causes an eastward migration of the convective zone, thus bringing rains to the east, droughts to the west. The east–west Intertropical Convergence Zone, simultaneously moves equatorward. During La Ninã, a westward contraction of the warm waters shifts the convection zone back westward and intensifies the trades. The Southern Oscillation is this interannual back-and-forth movement of air masses across the tropical Pacific. Figure 2 depicts its complementary El Ninõ and La Ninã states. A hierarchy of models is available to simulate the atmospheric response to changes in tropical sea surface temperature patterns. The most realistic, the general circulation models used for weather prediction, attempt to incorporate all the important physical processes that determine the atmospheric circulation. Simpler models isolate a few specific processes and, by elucidating their roles, provide physical insight. Particularly useful for capturing the essence of the Southern Oscillation – the departure from the time-averaged state – is a model that treats the atmosphere as a one-layer fluid on a rotating

sphere. Motion is driven either by a heat source over the region of unusually high surface temperatures (the moisture carried by the winds that converge onto the source can amplify the magnitude of the heat source), or is driven by sea surface temperature gradients that give rise to pressure gradients in the lower layer of the atmosphere. Such models are widely used as the atmospheric components of relatively simple coupled ocean–atmosphere models. The complex general circulation models reproduce practically all aspects of the interannual Southern Oscillation over several decades if, in the simulations, the observed sea surface temperature patterns are specified. The models fail to do so if the temperature patterns are allowed only seasonal, not interannual variability. This important result implies that, although weather prediction is limited to a matter of weeks, coupled ocean–atmosphere models should be capable of extended forecasts of certain averaged fields, those associated with the Southern Oscillation, for example. The key difference between weather and climate is that the first is an initial value problem – a forecast of the weather tomorrow requires an accurate description of the atmosphere – whereas climate (from an atmospheric perspective) is a boundary value problem; a change in climate can be induced by altering conditions at the lower boundary of the atmosphere. The crucial conditions are sea surface temperatures in the case of El Ninõ.

The Oceans The salient feature of the thermal structure of the tropical oceans is the thermocline, the thin layer of large vertical temperature gradients at a depth of approximately 100 m, that separates warm surface waters from colder water at depth. Sea surface temperature patterns in the tropics tend to reflect variations in the depth of the thermocline and those variations are controlled by the winds. Thus the surface waters are cold where the thermocline is shallow, in the eastern equatorial Pacific for example, and are warm where the thermocline is deep, in the western tropical Pacific. The downward slope of the thermocline, from east to west in the equatorial Pacific, is a consequence of the westward trade winds that, along the equator where the Coriolis force vanishes, drive the surface waters westward. When those winds relax, during El Ninõ, the warm surface waters in the west return to the east so that the thermocline there deepens and sea surface temperatures rise. To a first approximation, the change from La Ninã to El Ninõ amounts to a horizontal (east–west), adiabatic redistribution of warm surface waters, and is the dynamical response of the ocean to the changes in the winds.



243

~ conditions La Nina

North-east

180°

Trade winds

Equator Very intense S.E. trade winds Strong high pressure zone

25°C 20°C

Steeply sloping ther

mocline

~ conditions EI Nino

28°C

Equator Weak trade winds Westerly winds 180°

Weak high pressure zone

25°C 20°C

Thermocline that is almost hor izontal

Figure 2 A schematic view of La Ninã (top) and El Ninõ (bottom) conditions. During La Ninã, intense trade winds cause the thermocline to have a pronounced slope, down to the west, so that the equatorial Pacific is cold in the east, warm in the west where moist air rises into cumulus towers. The air subsides in the east, a region of little rainfall, except in the doldrums where the south-east and north-east trades converge. During El Ninõ, the trades along the equator relax, as does the slope of the thermocline when the warm surface waters flow eastward. The change in surface temperatures is associated with an eastward shift of the region of heavy precipitation.

Theoretical studies of the oceanic response to changes in the wind started with studies of the generation of the Gulf Stream. How long, after the sudden onset of the winds, before a Gulf Stream appears and the ocean reaches a state of equilibrium? The answer to this question provides an estimate of the ‘memory’, or adjustment-time of the ocean, a timescale that turns out to be relevant to the timescale of the Southern Oscillation. In an unbounded ocean,

a Gulf Stream is impossible, so that the generation of such a current depends on the time it takes for information concerning the presence of coasts to propagate east–west across the ocean basin. The speed of the oceanic waves that carry this information, known as Rossby waves, increases with decreasing latitude. It should take on the order of a decade to generate the Gulf Stream from a state of rest, a matter of months to generate the same current


244


near the equator. Fortuitously, the three tropical oceans have similarities and differences that provide a wealth of information about the oceanic response to different wind stress patterns. To explain and simulate tropical phenomena, oceanographers developed a hierarchy of models of which the simplest is the shallow-water model whose free surface is a good analog of the very sharp, shallow tropical thermocline. Studies with that tool showed how a change in the winds over one part of an ocean basin (e.g. the western equatorial Pacific), can influence oceanic conditions in a different and remote part of the basin (e.g. off the coast of Peru). Thus a warming of the surface waters along the coast of Peru during El Ninõ could be a consequence of a change in the winds over the western equatorial Pacific. The roles of waves and currents in effecting such changes depend on the manner in which the winds vary. If the winds change abruptly, then waves are explicitly present, but if the winds vary slowly, with a timescale that is long compared with the adjustment time of the ocean, then the waves are strictly implicit because the response is an equilibrium one. A detailed description of the response of the Pacific to slowly varying winds first became available for El Ninõ of 1982 and provided a stringent test for sophisticated general circulation models of the ocean. The success of such models in simulating the measurements bolstered confidence in the models which are capable of reproducing, deterministically, the oceanic aspects of the Southern Oscillation between El Ninõ and La Ninã over an extended period, provided that the surface winds are specified. Today such models serve, on a monthly operational basis, to interpolate measurements from a permanent array of instruments in the Pacific, thus providing a detailed description of current conditions in the Pacific Ocean, a description required as initial conditions for coupled ocean–atmosphere models that predict El Ninõ.

Interactions between the Ocean and Atmosphere Suppose that, during La Ninã, when intense trades drive the surface waters at the equator westward, a random disturbance causes a slight relaxation of the trades. Some of the warm water in the west then starts flowing eastward, thus decreasing the east– west temperature gradient that maintains the trades. The initial weakening of the winds is therefore reinforced, causing even more warm water to flow eastward. This tit-for-tat (positive feedback) leads to the demise of La Ninã, the rise of El Ninõ. The latter state can similarly be shown to be unstable to random perturbations.

A broad spectrum of natural modes of oscillation – the Southern Oscillation is but one – is possible because of unstable interactions between the ocean and atmosphere. The properties of the different modes depend mainly on the mechanisms that control sea surface temperature, because that is the only oceanic parameter that affects the atmosphere on the timescales of interest here. Those mechanisms include advection by oceanic currents, and vertical movements of the thermocline caused either by local winds, or by remote winds that excite waves that propagate along the thermocline. In one class of coupled ocean– atmosphere modes, sea surface temperature variations depend primarily on advection. Consider an initial perturbation in the form of a confined equatorial region of unusually high sea surface temperatures. That warm patch, superimposed on waters that get progressively warmer from west to east, induces westerly winds to its west, easterly winds to its east. The westerly winds drive convergent currents that advect warm water, in effect extending the patch westward. The easterly winds induce divergent currents that advect cold water, thus lowering sea surface temperatures and contracting the warm patch on its eastern side. The net result is a westward displacement of the original disturbance. A mode of this type is involved in the response of the eastern equatorial Pacific to the seasonal variations in solar radiation. In certain coupled ocean–atmosphere modes the slow adjustment of the ocean (over a period of months and longer) to a change in winds, in contrast to the short period (weeks) the atmosphere takes to come into equilibrium with altered sea surface temperature, is of great importance. (That is why the ocean, far more than the atmosphere, needs to be monitored to anticipate future developments.) These modes, which are known as ‘delayed oscillator’ modes because of the lagged response of the oceans, differ from those discussed in the previous paragraph, in being affected by the boundedness of the ocean basin – the presence of north–south coasts. Furthermore, sea surface temperatures now depend mainly on thermocline displacements, not on upwelling induced by local winds. The effect of winds to the west of an, initially, equatorially confined region of unusually warm surface waters can extend far eastward because of certain waves that travel efficiently in that direction along the equator. This type of mode is associated with long timescales, of several years, on which the oceanic response is almost, but not quite, in equilibrium with the gradually changing winds. That distinction is of vital importance because the small departure from equilibrium, the ‘memory’ of the system, brings about the transition from one phase of the oscillation to the next. If El Ninõ conditions are in



existence then the delayed response of the ocean causes El Ninõ to start waning after a while, thus setting the stage for La Ninã. To which of the various modes does the observed Southern Oscillation correspond? The answer depends on the period under consideration, because the properties of the dominant mode at a given time (the one likely to be observed) depend on the background state at the time. For example, if the thermocline is too deep then the winds may be unable to bring cold water to the surface so that ocean–atmosphere interactions, and a Southern Oscillation, are impossible. Such a state of affairs can be countered by sufficiently intense easterly winds along the equator because they slope the thermocline down towards the west, shoaling it in the east. These considerations indicate that the time-averaged depth of the thermocline, and the intensity of the easterly winds, are factors that determine the properties of the observed Southern Oscillation. (The temperature difference across the thermocline is another factor.) If the thermocline is shallow, and the winds are intense (this was apparently the case some 20 000 years ago during the last Ice Age), then the Southern Oscillation is likely to have a short period of 1 or 2 years, and to resemble the mode excited by the seasonal variation in solar radiation. In general, the observed mode is likely to be a hybrid one with properties intermediate between those associated with the seasonal cycle, and those known as the delayed oscillator type. During the 1960s and 1970s the thermocline was sufficiently shallow, and the trade winds sufficiently intense, for the dominant mode to have some of the characteristics of the annual mode – a relatively high frequency of approximately 3 years, and westward phase propagation. Since the late 1970s, the background state has changed because of a slight relaxation of the trades, a deepening of the thermocline in the eastern tropical, and a rise in the surface temperatures of that region. This has contributed to a change in the properties of El Ninõ. During the 1980s and 1990s it was often associated with eastward phase propagation, and with a longer period of recurrence, 5 years. This gradual modulation of the Southern Oscillation cannot explain the differences between one El Ninõ and the next, why El Ninõ was very weak in 1992, exceptionally intense in 1997. Those differences are attributable to random disturbances, such as westerly wind bursts that last for a week or two over the far western equatorial Pacific. They can be influential because the background state, for the past few decades, has always been such that interactions between the ocean and atmosphere are marginally unstable at most, or slightly damped so that the continual Southern Oscillation is sustained by

245

sporadic perturbations. A useful analogy is a swinging pendulum that is subject to modest blows at random times. A blow, depending on its timing, can either amplify the oscillation – that apparently happened in 1997 when a burst of westerly winds in March of that year led to a very intense El Ninõ – or can damp the oscillation. The predictability of El Ninõ is therefore limited, because the westerly wind bursts cannot be anticipated far in advance. (Several models predicted that El Ninõ would occur in 1997, but none anticipated its large amplitude.) These results concerning the stability properties of ocean–atmosphere modes come from analyses of relatively simple coupled models: the ocean is a shallow-water model with a mixed, surface layer in which sea surface temperatures have horizontal variations; the atmosphere is also a single layer of fluid driven by heat sources proportional to sea surface temperature variations. These models deal only with modest departures from a specified background state that can be altered to explore various possible worlds. The results are very helpful in the development of more sophisticated coupled general circulation models of the ocean and atmosphere which have to simulate, not only the interannual Southern Oscillation, but also the background state. At this time the models are capable of simulating with encouraging realism various aspects of the Earth’s climate, and of the Southern Oscillation. As yet, the properties of the simulated oscillations do not coincide with those of the Southern Oscillation as observed during the 1980s and 1990s, presumably because the background state has inaccuracies. The models are improving rapidly.

Conclusions The Southern Oscillation, between complementary El Ninõ and La Ninã states, results from interactions between the tropical oceans and atmosphere. The detailed properties of the oscillation (e.g. its period and spatial structure) depend on long-term averaged background conditions and hence change gradually with time as those conditions change. The ocean, because its inertia exceeds that of the atmosphere by a large factor (oceanic adjustment to a change in forcing is far more gradual than atmospheric adjustment), needs to be monitored in order to anticipate future developments. The development of computer models capable of predicting El Ninõ is advancing rapidly.

See also Elemental Distribution: Overview. El Ninõ Southern Oscillation (ENSO). Rossby Waves. Satellite Remote


246


Sensing of Sea Surface Temperatures. Wave Generation by Wind.

Further Reading The Journal of Geophysical Research Volume 103 (1998) is devoted to a series of excellent and detailed reviews of various aspects of El Ninõ and La Ninã. Neelin JD, Latif M, and Jin F-F (1994) Dynamics of coupled ocean–atmosphere models: the tropical problem. Annual Review of Fluid Mechanics 26: 617--659.

Philander SGH (1990) El Ninõ, La Ninã and the Southern Oscillation. New York: Academic Press. Zebiak S and Cane M (1987) A model El Ninõ Southern Oscillation. Monthly Weather Review 115: 2262--2278. Schopf PS and Suarez MJ (1988) Vacillations in a coupled ocean–atmosphere model. Journal of Atmospheric Science 45: 549--566.


ELECTRICAL PROPERTIES OF SEA WATER R. D. Prien, Southampton Oceanography Centre, Southampton, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 832–839, & 2001, Elsevier Ltd.

Introduction The oceans are the biggest reservoir of electrolyte solution on Earth. The electrical properties of the sea water are mainly determined by the ions dissolved in it when DC or low-frequency AC electric fields are involved. It should be noted, however, that the water plays an important role in producing the ions from the electrically inert salt crystals. For high-frequency electric fields and electromagnetic wave propagation, the properties of the water itself become more important, since the mechanism of conduction changes. The electrical conductivity has become the most important electrical property of sea water because the definition of salinity is based on conductivity ratios. Thus, salinity in its current definition is an electrical property. Other consequences of the sea water being an electrolyte might not be as obviously relevant to the marine scientist as is the conductivity. Whenever instruments that consist of metal parts are lowered into the sea, the galvanic potential has to be considered, either because it might influence electric measurements carried out by the instrument or because of the accelerated corrosion it causes. The fact that the ions in the water are moved with it in the Earth’s magnetic field by the ocean currents can be used to determine water velocities. Finally, the optical properties of the sea water are by their nature electromagnetic properties.

example, has a lower conductance than a thick, short cylinder of the same volume. To get a measure independent of the shape and size of the body of a material, the conductivity of media is given in terms of the specific conductivity s. For a cylindrical volume of material with length l, cross-sectional area A, and resistance R between its plane, parallel faces, the specific conductivity is given by s¼

Source of Ions

While pure fresh water conducts electric current only very weakly (owing to a slight dissociation of the water molecules), the addition of salt (e.g., NaCl) to the water results in a significant rise in conductivity, although the compound NaCl is electrically neutral. It is known, however, that salt crystals consist of positive and negative charged ions that are attracted together by the Coulomb force given by F¼

Electrical conductance is the property of a body quantifying the ability to conduct an electric current. The electrical conductance G is given as the ratio of electric current I and voltage V and therefore equals the reciprocal of electrical resistance R: I 1 ¼ V R

Q1 Q2 4pe0 er2

½3

where Q1 and Q2 are the charges of the ions, e0 is the permittivity of vacuum, e is the permittivity of the medium, and r is the distance between the ions. To separate the ions, an amount of work

W¼

½1

The electrical conductance of a body is dependent not only on its material but also on the geometric properties of the body. A thin, long cylinder, for

½2

Thus, the specific conductivity is normalized by the geometry and is a property of the material only. When considering the electrical conductivity of a material one must differentiate between direct-current and low-frequency conductivity on the one hand and high-frequency conductivity on the other. The latter will be discussed in a separate section below. Whereas in metals the electric conductivity is based on electron conductance only, in electrolytic fluids –and therefore in sea water –the conductivity is based on ion conductance only, as far as direct-current and low-frequency conductivity is concerned. Both types of ions, positively and negatively charged, contribute to the electrical conductivity.

Electrical Conductivity

G¼

l 1 1 O m ¼ S m1 AR

ZN

F dr

½4

b

has to be carried out, where b is the distance between the Naþ and the Cl ions in the salt crystal (for NaCl, bC0.2 nm). In air (eC1) the work to


247

248

ELECTRICAL PROPERTIES OF SEA WATER

dissociate a NaCl molecule is 1.15 1018 J, but in water (eC81) only 1.42 1020 J is needed. Owing to their polar nature, the water molecules in the vicinity of the ions are orientated according to the charge of the ions and form a layer around the ions. This process is called hydratization and it sets energy free that then is used to separate the ions of the NaCl molecules. Thus, after salt is added to the pure water the dissolved salt molecules are dissociated into ions. The process of separation of the ions of the dissolved molecules is called electrolytic dissociation. The fact that the salt molecules are dissociated can be seen not only by carrying out a conductivity measurement (which implies the application of an electric field between the measuring electrodes) but also from the properties of the electrolyte. It is known that in solutions an osmotic pressure is present that leads to a decrease of vapor pressure and ice point as well as an increase of the boiling point. These effects are proportional to the concentration of the solution, as is shown by the Raoult laws. The solutions of water, however, show deviations of these properties from the values predicted by the Raoult laws. This can be explained only by the dissociation of the solute, because its effective concentration is higher. The effective concentration can be obtained by combining the concentrations of the dissociated components. Therefore, the effects mentioned above are more pronounced for electrolytic solutions. It seems surprising at first that, for example, the sodium ions can exist in the water without the wellknown reaction 2Na þ 2H2 O-2NaOH þ H2 m taking place but it has to be borne in mind that the chemical properties of atoms and molecules are determined almost exclusively by the electron configuration of the outer electron shell, which in the case of the Naþ ion is a noble gas configuration and therefore quite different from the configuration of a sodium atom. Mechanism of Conductivity in Electrolytes

A potential difference between two electrodes that are lowered into an electrolyte creates an electric field in the electrolyte, by which positively charged ions (anions) are forced towards the cathode and negatively charged ions (cations) towards the anode. At the electrodes the anions receive electrons from the cathode while the cations donate electrons to the anode. The movement of ions in the water therefore results in a flow of electrons between the electrodes and thus a current in the external circuit. It should be

_

+

=

E _

+

Cl

Na +

Na

_

_

Na +

Na +

Na

+

Na +

Cl

Cl e

_

Cl

_

+

Cl

_

Na +

Cl

Na

+

Na

+

Na

_

Cl

_

Cl

_e

Cl

_

Cl

Figure 1 Schematic of ionic conductivity mechanism in sea water. The ions are moving towards the electrodes under the influence of the electric field created by the voltage applied between the electrodes. When the ions reach the electrodes, cations donate electrons while anions receive electrons. Thus, an electric current between the electrodes results.

noted that current flow in electrolytic fluids (in contrast to current flow in metals) is connected to a mass flow (Figure 1). The electric field created by the voltage applied to the electrodes results in a force on the ions, which are accelerated towards the electrodes. Since the ions are moving in the fluid, friction opposes the electric force and after a short period of acceleration the frictional force balances the electric force and the ions move with constant velocity. According to Stokes’ law, the frictional force and therefore the resulting velocity is a function of the size of the ions. It should be mentioned that the friction of the ions in the fluid results in an increase of temperature of the fluid. The water molecules do not play an important role in the low-frequency or DC conductivity of the electrolyte. However, the water molecules have a dipole moment because of their polar nature. Each hydrogen–oxygen bond is polar covalent, with the hydrogen end positive with respect to the oxygen end. When no external electric field is applied, thermal agitation keeps the molecular dipoles randomly oriented; when a field is applied, the dipoles align themselves and the fluid takes on an electric polarization.

Factors Determining Conductivity

The conductivity of an electrolyte is a function of the concentration of ions. The greater the number of ions that are moving towards the electrodes, the greater is the number of electrons that can be interchanged at the electrodes and the higher is the resulting current. If all molecules of the solute are dissociated, the conductivity is a linear function of the concentration



_1

Conductivity, (S m )

of the solute. This linear relation between concentration of the solute and conductivity of the electrolyte holds true only for low concentrations. It was shown above that the dissociation of the salt molecules depends on the permittivity of the water. The permittivity is a macroscopic property and if a salt molecule is surrounded not only by water molecules but also by other salt molecules, the permittivity in the vicinity of the salt molecule is altered. At higher concentrations not all of the salt molecules are dissociated and therefore the increase of conductivity of the electrolyte as a function of concentration is weakened. This effect can be seen in Figure 2, which shows the conductivity as a function of salinity and temperature. The conductivity is not linearly increasing with increasing salinity; the slope is decreasing for higher salinities. For high-precision determination of salinity this effect has to be taken into account, although the salt concentration in sea water is not very high. Another factor determining the conductivity of an electrolyte is the mobility of the ions. As was mentioned above, the size of the ions determines their velocity in the electrolyte when under the influence of an electric field. The higher this velocity, the more ions reach the electrode in a given time interval and thus the higher is the current. The conductivity of an electrolyte is also a function of temperature. The viscosity of the solvent (here, water) decreases with temperature and the ion mobility is increased. The degree of dissociation of the electrolyte also changes with temperature. Some electrolytes show an increase of the degree of dissociation with higher temperatures; most of them,

8 6 4 2 0 40

30 30

Salin

20 20

ity, S

10

10 0

re,

atu per

) t (°C

Tem

Figure 2 Conductivity of sea water as a function of temperature and salinity. The conductivity was calculated assuming an absolute value of conductivity s(S ¼ 35, t ¼ 151C, p ¼ 0 dbar) ¼ 4.2194 S m1. The bold lines denote constant conductivities from 1 S m1 to 7 S m1.

249

however, show a decrease of the degree of dissociation with an increase of temperature. This explains why for some solutions the conductivity first increases with temperature and then decreases, after passing a maximum of conductivity, for further temperature increase. Definition of Salinity of Sea Water

Since the end of the nineteenth century it has been known that the composition of sea water is almost constant in space and time. It is therefore justified, to a good first-order approximation, to assume that sea water consists of just two components, the first being pure water and the second representing all dissolved ions that contribute to the mass of sea water. The concept of constant composition allows the choice of one type of ions to represent all other types. In the early 1900s, the salinity S of sea water was defined in terms of chlorinity Cl, defined as the chlorideequivalent mass ratio of halides to the mass of sea water, by a linear least-squares regression formula fitted to mass ratios of salt residues of nine sea water samples evaporated to dryness. The regression formula, known as the Knudsen formula, is S ¼ 0:030 þ 1:8050 Cl

½5

In the mid to late 1950s, the use of conductivity bridges for salinity determination was increasing and led to a concern about using a chlorinity standard to serve as a conductivity standard. In 1960, a study of the relationships between chlorinity, conductivity, and density on several hundred samples of sea water from widely distributed locations was undertaken at the British National Institute of Oceanography. A markedly higher correlation was found between density and conductivity than between density and chlorinity, and subsequently a Joint Panel on the Equation of State of Seawater was formed to review the relationship of the equation of state to chlorinity, salinity, electrical conductivity, and refractive index. The findings of the panel led to the definition of salinity in terms of the electrical conductivity ratio R15 at 151C to that of sea water having a salinity of 35. The initial recommendation was to define salinity in terms of density, which would have allowed construction of formulas relating density to a measured variable such as electrical conductivity, refractive index, or density itself. It turned out, however, that the necessary precision for the determination of absolute density and conductivity of standard sea water was not achievable at the time.


250


Owing to problems arising from using standard sea water as a reference, it was agreed to use a potassium chloride solution as a standard to determine the salinity of the standard sea water. This definition was published as the Practical Salinity Scale 1978 (PSS78). Therefore, the salinity of sea water in its current definition is an electrical property. In the following the nomenclature of the Unesco technical papers in marine science will be used, where C denotes the conductivity and R the conductivity ratio. If CðS; t; pÞ is the electrical conductivity of sea water at practical salinity S (PSS78), temperature t (International Practical Temperature Scale 1968 (IPTS68)) and pressure p in decibars, the conductivity ratio is defined to be

with the coefficients e1 ¼ þ2:070 105 e2 ¼ 6:370 1010 e3 ¼ þ3:989 1015 d1 ¼ þ3:426 102 d2 ¼ þ4:464 104 d3 ¼ þ4:215 101 d4 ¼ 3:107 103 The ratios rt can be calculated from the temperature as rt ¼ c0 þ c1 t þ c2 t2 þ c3 t3 þ c4 t4

CðS; t; pÞ R¼ Cð35; 15; 0Þ

where

where C(35, 15, 0) is the conductivity of standard sea water of practical salinity 35, and 151C and atmospheric pressure, defined to be equal to the conductivity of a reference solution of potassium chloride (KCl) at the same temperature and pressure. This KCl reference contains 32.4356 g of KCl in a mass of one kilogram of solution. The conductivity ratio can be factored into three parts: R ¼ Rp Rt rt

c0 ¼ þ6:766097 101 c1 ¼ þ2:00564 102 c2 ¼ þ1:104259 104 c3 ¼ 6:9698 107 c4 ¼ þ1:0031 109 The ratio Rt then can be obtained from

½7 Rt ¼

where Rp ðS; t; pÞ ¼

R Rp rt

With the ratio Rt and the temperature t, the practical salinity S then can be computed using the equation

CðS; t; pÞ CðS; t; 0Þ

1=2

S ¼ a 0 þ a1 R t Rt ðS; tÞ ¼

CðS; t; 0Þ Cð35; t; 0Þ

3=2

þ a2 Rt þ a3 Rt 5=2

þa5 Rt

þ DS

with the coefficients rt ðt Þ ¼

½9

½6

Cð35; t; 0Þ Cð35; 15; 0Þ

a0 ¼ þ0:0080

Since the ratios Rt or R are the properties measured with conductivity bridges together with temperature and pressure, formulas have been developed for the ratios in these variables rather than in salinity S. When measuring the conductivity ratio R together with temperature t and pressure p, the ratio Rp can be obtained as p e1 þ e2 p þ e3 p2 Rp ¼ 1 þ 1 þ d1 t þ d2 t2 þ ðd3 þ d4 tÞR

a1 ¼ 0:1692 a2 ¼ þ25:3851 a3 ¼ þ14:0941 a4 ¼ 7:0261 a5 ¼ þ2:7081 where

½8


X i

ai ¼ 35:0000

þ a4 R2t

½10


The term DS in equation[10] is given by t 15 DS ¼ 1 þ kðt 15Þ 1=2 3=2 5=2 b0 þ b1 Rt þ b2 Rt þ b3 Rt þ b4 R2t þ b5 Rt ½11 where b0 ¼ þ0:0005 b1 ¼ 0:0056 b2 ¼ 0:0066 b3 ¼ 0:0375 b4 ¼ þ0:0636 b5 ¼ 0:0144 P i bi ¼ 0:0000 and k ¼ þ0:0162 The given equations are valid over a temperature range from 21C to 351C and practical salinity from 2 to 42. The pressure range is not explicitly stated but a table of maxima of the range of pressures over which the pressure correction to conductivity ratio was computed is given in the original publication. It should be noted that no absolute conductivity value is given in the PSS78 since it is not required. A reference value given in literature is C(35, 15, 0) ¼ 4.2194 S m1.

251

which means that the molecules undergo rapid rotations, aligning themselves with the electric field. The sea water therefore takes on an orientational polarization. The conductivity and the permittivity become functions of the frequency of the electromagnetic field. When the frequency is too high for the molecules to be able to follow the field changes, because of their inertia, their contribution to the polarization decreases and the permittivity drops markedly. For water, the permittivity is fairly constant (B81) for frequencies up to about 10 GHz, after which it falls off quite rapidly. At higher frequency, the electronic polarization becomes the main factor determining the permittivity. Electronic polarization arises when the electron clouds are shifted relative to the nuclei and generate a dipole moment. This means that for high-frequency electric fields the mechanism of conductivity changes and is no longer connected to a mass transport. While ion mobility is a main factor for low-frequency and DC conductivity, the dipole moment of the sea water and its frequency dependence play the major role in highfrequency conductivity. The change in mechanism also needs another formulation and the theory of electromagnetic wave propagation in matter must be applied. For isotropic homogeneous media and in the case of pure periodic functions of time of the form exp{iot}, where o is the frequency of the field, the Maxwell equations for the components of the magnetic field B and electric field E are reduced to the wave equations: DB þ q2 B ¼ 0

½12

DE þ q2 E ¼ 0

½13

Electromagnetic Wave Propagation The mechanisms of conductivity change dramatically when the voltage applied to two electrodes in sea water changes from DC or low frequency to high frequency. At low frequencies there is not much change in the mechanism of conductivity as the ions are accelerated up to their constant velocity and slow down again when the sign of the electric field changes. The conductivity is slightly decreased compared to the DC values because of repetitive acceleration of the ions. When the frequency is further increased, the ions are no longer accelerated to the point of constant velocity in the fluid, but are only slightly shifted from their position before the sign of the electric field changes and the ions are forced back. At even higher frequencies, the ions are not accelerated any more and the ion and its associated water molecules only change orientation according to the electric field and their electric dipole moment. The water molecules are also polarized according to their dipole moment,

Here D is the Laplace operator and the complex wavenumber q is related to the wavenumber in vacuo, q0, and the complex refractive index, m, by q ¼ mq0

½14

with q0 ¼

2p w ¼ l0 c

½15

and rffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi 4ps m ¼ em i o

½16

Here l0 is the wavelength in vacuum, c is the velocity of light in vacuum, e is the permittivity, m is the


252


magnetic permeability, and s is the conductivity of the medium. The complex refractive index m can be separated into the real and imaginary part: m ¼ n ik

½17

where n is the ordinary refractive index and k is the electrodynamic absorption coefficient. The permittivity e also is a complex quantity and can be written as e ¼ e0 ie00

½18

With mE1; an approximation valid for most materials, the two equations [16] and [17] for the refractive index m show the relation n and k to the permittivity and the conductivity: e0 ¼ n2 k2

½19

and e00 þ

4ps ¼ 2nk o

½20

The solutions of the wave equations are of the form exp½iðot qzÞ ¼ exp½kq0 zexp½iðot nq0 zÞ

½21

This is a damped plane wave propagating in the zdirection with a damping factor of kq0. The real and imaginary parts of the refractive index therefore represent two different properties, the ordinary refractive index n gives the phase advance of the wave, while the electrodynamic absorption coefficient k determines the reduction in amplitude of the wave. The dependence of n and k on the frequency of the field is governed by the various polarization mechanisms contributing to the dipole moment at the particular frequency.

Galvanic Couples When a metal body is lowered into water, ions of the metal will be dissolved in the water. An equilibrium is reached when the osmotic pressure of the metal ions dissolved in the liquid balances the metal’s solution tension. Since the metal electrode loses positively charged ions, it becomes negatively charged while the liquid becomes positively charged. Therefore, an electric field is created between the liquid and the electrode. The resulting electric force opposes the further solution of metal ions in the liquid and the system is in equilibrium. The potential difference between the metal and the water is called the

galvanic potential. It should be mentioned that the galvanic potential cannot be measured directly, since this would always involve a second electrode being lowered into the solution; therefore only the potential difference between two electrodes can be measured. It is this potential difference between two electrodes in the electrolyte that can lead to oxidation of one of the electrodes. If the electrodes are of different metals, they are called a galvanic couple. It should be noted that not only different metals form galvanic couples; other materials (e.g., carbon) can also be part of a galvanic couple. The galvanic potential between the electrolyte and the electrode is built up at both electrodes. Since the electrodes consist of different metals, their respective galvanic potentials are different and a voltage is built up between the electrodes. This voltage results in a current through the electrolyte and therefore in reduction of one electrode and oxidation of the other. The potential difference between two materials can be obtained from the electrochemical series, in which the potential differences of different materials in respect to a common reference electrode (e.g., the standard hydrogen electrode) are listed. The voltage between two electrodes of different materials is the difference of the listed values of the potential series. The electrochemical reduction and oxidation effects have to be considered when designing instruments for the marine environment. The housings of instruments have to be designed to exhibit as little oxidation in the sea water as possible. When different metals are used in the vicinity of each other (e.g., an instrument and a sensor housing), one of the metals is reduced. To avoid this process, electrodes are sometimes added to the instrument housing, either made of a suitable metal or connected to a voltage, so that all other metals of the instrument are at a negative potential to the electrolyte. The additional electrodes are dissolved slowly in the sea water and therefore have to be replaced from time to time. They are accordingly called sacrificial electrodes. Another method for avoiding corrosion by electrochemical processes is the surface coating of housings by adding an electrical isolation layer to the metal. A common treatment for aluminum, for example, is anodizing. The material to be treated is used as the anode for the electrolysis of a suitable electrolyte (e.g., sulfuric acid). An oxide layer is formed on the anode surface that is electrically isolating and can be hardened by further treatment to increase scratch resistance of the material.



Velocity Measurement The conductivity of the sea water is not the only electric property used for making measurements of the state of sea water. The fact that sea water is an electrolytic solution and contains free charges (the ions) is also used to determine current velocities on various scales. If a conductor is moved in a magnetic field, a current is induced in the conductor according to Faraday’s law. Since sea water is a conductor, a flow of sea water can be seen as the movement of a conductor and therefore Faraday induction can be used to measure the water velocity. If displacement currents in the water can be neglected, the electric potential distribution is given by the equation

proportional to the water flow perpendicular to the cable and the earth’s magnetic field. Measurements with Artificial Magnetic Fields

Another kind of electromagnetic velocity sensors uses the creation of magnetic fields by the use of coils. An alternating current in the coil induces an oscillating magnetic field. Two sensing electrodes are used to define the length of the conductor and the amplitude of the oscillating voltage between the electrodes is proportional to the velocity of the water. A constantvoltage signal can arise from galvanic potentials and therefore is subtracted. Thus, the water velocity can be determined from the oscillating voltage.

Symbols used

r 2 F ¼ r ðv B Þ

½22

where F is the electric potential, v is the velocity of the water and B is the magnetic field. In an ideal case, where the flow is laminar and uniform, the voltage V between two electrodes in contact with the sea water is V ¼ rBvn where r is the distance between the electrodes, B is the magnitude of the magnetic field, and vn is the velocity component normal to B and v. These velocity measurements can be carried out in the earth’s magnetic field or by generating a local magnetic field. Measurements in the Earth’s Magnetic Field

The geomagnetic electrokinetograph (GEK) consists of two electrodes that are towed through the water, behind a ship. The resulting voltage between the electrodes gives a measure of the vertical integrated flow perpendicular to the ship’s path. The expendable current profiler (XCP) is a sensor with two pairs of electrodes attached to the surface of the sensor, which sinks in the earth’s magnetic field. The resulting voltage between the electrodes measures the current velocity perpendicular to the sensor’s movement; it therefore measures the horizontal components of the water velocity. On descent, the probe is rotated around the vertical axis (the axis of descent) to determine any constant-voltage offsets induced by galvanic potentials. Another method for estimating large-scale, slowly varying currents is the use of seafloor cables. A crossbasin cable is used to sense the potential difference between the two sides of the basin, which is

253

Electrical conductivity G Conductance (S ¼ O 1) I Current (A) R Resistance (O) s Conductivity (S m 1) F Force (N) Q Charge (C) Permittivity of vacuum (As V 1 m 1) e0 e Relative permanittivity W Work (Nm) Definition of salinity of seawater S Salinity R Conductivity ratio C Conductivity (S m 1) t Temperature (1C) p Pressure (dbar) Electromagnetic wave propagation B Magnetic field (V s m 1) E Electric field (V s m 1) q Wavenumber m 1 l0 Wavelength (m) o Frequency (s 1) c Speed of light (m S 1) s Conductivity (s m 1) e Relative permittivity m Relative magnetic permeability n Refractive index k Electrodynamic absorption coefficient t Time (s) Velocity F v B V r

measurement Electric potential (V) Water velocity (m s ) Magnetic field (V s m ) Voltage (V) Distance (m)


254


See also CTD (Conductivity, Temperature, Depth) Profiler. Expendable Sensors. Inherent Optical Properties and Irradiance. Single Point Current Meters.

Further Reading Dexter SC (1979) Handbook of Oceanographic Engineering Materials, Ocean Engineering Series. New York: Wiley-Interscience. Fofonoff NP (1985) Physical properties of seawater: a new salinity scale and equation of state for seawater. Journal of Geophysical Research 90(C2): 3332--3342. Krause G (1986) In-situ instruments and measuring techniques. In: Su¨ndermann J (ed.) Landolt-Bo¨rnstein, New

Series, Group V, Geophysics vol. 3, Oceanography, Subvolume a, pp. 134–232. Berlin: Springer. Robinson RA and Stokes RH (1959) Electrolyte Solutions. London: Butterworth. Sanford TB (1971) Motionally induced electric and magnetic fields in the sea. Journal of Geophysical Research 76: 3476--3492. Shifrin KS (1988) Physical Optics of Ocean Water, AIP Translation Series. New York: American Institute of Physics. Unesco (1983) Algorithms for computation of fundamental properties of seawater. Unesco technical papers in marine science, 44, p. 53. Paris: Unesco. von Arx WS (1950) An electromagnetic method for measuring the velocities of ocean currents from a ship under way. Papers in Physical Oceanography and Meteorology 11(3): 1--62.


ELEMENTAL DISTRIBUTION: OVERVIEW Y. Nozaki, University of Tokyo, Tokyo, Japan Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 840–845, & 2001, Elsevier Ltd.

Introduction More than 97% of liquid water on the earth exists in the ocean. The ocean water contains approximately 3.5% by weight of dissolved salt. What is the elemental composition of the salts, how does it vary from place to place and with depth, and why? These are fundamental questions for which chemical oceanographers have sought answers. Despite more than a hundred years of intense investigation by modern chemical oceanography, the answers have not been fully elucidated. Nevertheless, we are now approaching complete understanding of the chemical composition of sea water and its variability in the ocean.

Historical Review By the late nineteenth century it was well-established that the major components of sea water are extremely constant in their relative abundance, and comprise some ten constituents including Cl, Naþ, Mg2þ, SO2 4 (see Conservative Elements). The analytical results reported by W. Dittmar in 1884 for waters collected during the British RMS Challenger Expedition (1872–1876) from the world’s oceans were almost the same as today’s values. The constancy of major chemical composition has led oceanographers to define ‘salinity’ as a fundamental property together with temperature to calculate the density of sea water. It was routine for classic physical oceanographers to titrate sea water for chloride (plus bromide) ion with silver nitrate standard solution, until the mid 1960s when salinity could be determined more practically by measurement of conductivity. On the other hand, for minor elements, there has been little information gained since the establishment of major chemical composition of sea water. Measurements of trace constituents in sea water are difficult because of their very low abundance. There was a clear tendency for the reported concentrations of many trace elements to become lower and lower as time elapsed. This trend was, of course, not real but an artifact. It is a famous story that, to aid

Germany’s national deficit after World War I, the Nobel Prize winning chemist F. Haber attempted to recover gold from sea water which according to the current literature occurred at about 5mg m3. He completely failed however, but, after long and rigorous examination, he found that the concentration was B1000 times less than that expected. Incidentally, Haber’s value of gold concentration was two orders of magnitude higher compared to later reports (Table 1). Another good example may be found in the measurements of lead in sea water by Patterson and his associates (see Anthropogenic Trace Elements in the Ocean). The vertical profile of Pb in the North Pacific obtained by Schaule and Patterson in 1981 had concentrations about two orders of magnitude lower than those reported earlier (B1970) by the same workers although their 1981 values are believed to be accurate and real (Table 1).

Technical Challenge It is now known that the most obvious reason for these trends is the continuous improvement in removing sources of contamination during sampling, handling, storage, and analysis. Significant efforts and advances in such field and laboratory techniques had been made until the GEOSECS (Geochemical Ocean Section Study) program started at around 1970. For example, polyvinyl chloride Niskin-bottle multisampling system together with CTD (conductivity–temperature–depth) sensors has routinely been employed in the hydrocasts, replacing the serial Nansen (metallic) bottle sampling method most widely used prior to that time. Yet, this was not enough for many trace metals except for barium, and an intercalibration exercise made in the early stage of the program did not produce any congruent results between laboratories. It was a significant and wise decision of the GEOSECS leaders that they focused on radionuclides and stable isotopes, which are almost free from contamination, and did not get involved in trace element geochemistry. Obviously, without having the real concentration data, any arguments that might be built upon them would be meaningless. Obtaining clean (uncontaminated) water samples from various depths of the ocean is of prime importance in the study of trace metals. In this regard, various types of sampling bottles have been


255

256

ELEMENTAL DISTRIBUTION: OVERVIEW

Table 1

Estimated mean oceanic concentrations of the elements



developed both domestically and commercially. They include the Cal-Tech Patterson sampler, modified Go-Flo bottles, and lever-action or X-type Niskin bottles. None of them are easy to keep clean and handle properly, and experience is needed in their operation depending on the type of bottle. Hydrowire is also important, since normal steel wire has rust and grease that can easily contaminate the water. To avoid this, some workers use plastic Kevler line and others use stainless-steel wire or a titanium armored cable.

257

With the rapid growth of semiconductor industries from the early 1970s, clean laboratory techniques also become more popular in the field of marine chemistry and helped considerably to reduce contamination from reagents, containers, and dust in the room atmosphere. Real oceanic concentrations of trace metals are so low that conventional analytical techniques prior to 1970 were not normally sensitive enough to detect them except in polluted or some coastal waters. Thus, significant efforts were also devoted to developing more sensitive and reliable


258


methods using atomic absorption spectrophotometry, chemiluminesence detection, isotope dilution mass spectrometry, etc. As a result, in the late 1970s, data of some transition metals, like Cd, Cu, and Ni were obtained by the Massachusetts Institute of Technology group and soon after confirmed by others using different or modified methods. Their oceanic profiles were quite consistent with known biogeochemical cycling and scavenging processes in the ocean. Thus, these features have often been referred as ‘oceanographically or geochemically consistent’ distribution by subsequent workers. Since then, growing numbers of publications describing the oceanic distributions of trace elements in sea water based on modern technologies have appeared year by year.

Oceanic Profiles It is now possible to compile, with reasonable confidence, the vertical profiles in the form of Periodic chart (Figure 1), where the data from the North Pacific have been chosen since physical processes that affect the distribution are relatively simple and well documented. Figure 1 is an updated version of the original, including new data for Nb, Ta, Hf, Os, Ag, and rare earth elements. Now, there remains only one element, Ru on which no real data have been reported (see Platinum Group Elements and their Isotopes in the Ocean). However, confirmation is needed for many elements, including Sn, Hg, Rh, Pd, Au, Ir, Pt, etc., since they are based on a single study or on controversial results by different workers. Nevertheless, it is clear that the long-standing dream to establish the chemical composition of sea water is about to become a reality. Trace elements follow one or more of the categories which are described below. Conservative type Some of the trace elements such as U, W, and Re form stable ionic species, 2 UO2(CO3)2 2 , WO4 , and ReO4 in sea water. Hence, their oceanic behavior is conservative (follow salinity) and their mean residence times in the ocean are generally long (e.g., >105 years). There is no significant variation in their concentration between different oceanic basins. Recycled type (nutrient-like) Many others, e.g., Ni, Cd, Zn, Ge, and Ba, show a gradual increase in their concentration from the surface to deep water, much like nutrients (nitrate, phosphate, and silicate or alkalinity), suggesting their involvement in the biogeochemical cycle of biological uptake in the

surface water and regeneration in deep waters. As a result of global ocean circulation, the deep-water concentrations of this type are higher in the Pacific than in the Atlantic. Scavenged type Trace metals such as Al, Co, Ce, and Bi, show surface enrichment and depletion in deep waters, in contrast to the opposite trend in nutrient types. These elements are highly particlereactive and are rapidly removed from the water column by sinking particulate matter and/or by scavenging at the sediment–water interface. Their mean oceanic residence times are short (o102–103 years). Interoceanic variations in their concentration can be large (e.g., Atlantic/Pacific concentration ratioB40 for Al) depending on kinetic balance between supply and removal for the specific basins. Redox-controlled type Elements such as Cr, As, Se, and Te exist in sea water at more than one oxidation state. Their oceanic behavior is strongly dependent on the chemical form. Their reduced states are thermodynamically unstable in normal oxygenated waters but are probably formed through biological mediation. Reduced species can also be formed in anoxic basins, the Black Sea, Cariaco Trench, some fiords, and in organic-rich sediments. Anthropogenic and transient type Finally, Pb and Pu are good examples of elements whose oceanic distributions are globally influenced by human activities (see Anthropogenic Trace Elements in the Ocean). Their oceanic distributions are changing with time. Although some others, such as Hg, Sn, Cd, and Ag, are deduced to be similarly influenced, their transient nature has not yet been proven through direct observation.

Particle Association and Speciation One of the important features of Figure 1 is that the concentration, even for trace elements, varies fairly smoothly and continuously with depth. This casts doubt on some erratic and highly discontinuous values unless there are obvious reasons for them, such as hydrothermal influence or difference in the water masses. The data shown in Figure 1 are largely based on filtered samples and therefore, can be referred as ‘dissolved concentration.’ For conservative elements, it does not matter whether the water sample is filtered or not, since there is virtually no difference in the analytical results. For most nutrienttype elements, particle association in the open ocean


259

Figure 1 Vertical profiles of elements in the North Pacific Ocean.



260


is generally small (oB5%) and therefore, the gross features of unfiltered samples remains the same as dissolved samples. However, filtration becomes important for coastal waters, and certainly for scavenged-type elements in any place, since particle association could easily exceed dissolved concentration. Various types of membrane filters with different pore sizes, generally in the range B1–0.04 mm, have been used, and, therefore, the so-called ‘dissolved’ concentration includes different amounts of colloidal form. Furthermore, ionic species of many trace elements form complexes with ligands in sea water. Table 1 lists the most probable inorganic species as deduced from thermodynamics. Recently, complexation of heavy metals (e.g., Fe, Zn, Cu, and Co) with organic ligands which occur at nanomolar concentration levels in the surface water has been investigated (see Transition Metals and Heavy Metal Speciation). These organic complexes are particularly important for understanding the roles of trace elements in plankton biology and metabolism. However, the technologies to detect and separate these metal–organic complexes are not yet available.

to large contamination problems in various stages of sampling and analysis. It was only during the last quarter of the twentieth century that those problems were eventually overcome by current efforts of chemical oceanographers. Since then, more and more reliable data have accumulated and now the oceanic distribution is known for almost all the elements. Some of the trace metals, such as Al, and rare earth elements (see Transition Metals and Heavy Metal Speciation) serve as useful tracers of water masses in describing hydrographic structures and patterns of ocean circulation. Others, such as Fe, Si, and perhaps Zn, have an essential role in phytoplankton growth and hence affect the global carbon cycle through the ‘biological pump.’ Resemblance in the oceanic distribution between Cd and phosphorus, and between Ba and alkalinity, provides a basis by which those heavy metals can serve as useful proxies in reconstructing the paleo-oceanographic environment from deep-sea sediment strata.

See also Anthropogenic Trace Elements in the Ocean. Conservative Elements. Refractory Metals. Transition Metals and Heavy Metal Speciation.

Summary The chemical constituents of sea water show a very wide range in their concentration of more than 15 orders of magnitude, from chlorine (B0. 5 mol kg1) to the least abundant platinum group elements (e.g., IroB1 fmol kg1). The measurement of trace elements in sea water is extremely difficult for a long time and was not achieved properly for most elements due

Further Reading Nozaki Y (1997) A fresh look at element distribution in the North Pacific Ocean. EOS Transactions, AGU 78: 221. Li YH (1991) Distribution patterns of elements in the ocean: a synthesis. Geochimica et Cosmochimica Acta 55: 3223--3240.


ENERGETICS OF OCEAN MIXING A. C. Naveira Garabato, University of Southampton, Southampton, UK & 2009 Elsevier Ltd. All rights reserved.

Introduction One of the defining features of the ocean’s physical environment is its nearly ubiquitous stable stratification (Figure 1). Aside from a relatively thin and homogeneous layer that is widely found near the surface (the so-called upper ocean mixed layer), the density of the ocean increases monotonically with depth in a perceptible manner, the rate of this increase generally declining toward the ocean floor. Current views on the origin of the ocean stratification began to

take form in the early twentieth century, as the oceanographers Georg Wu¨st and Albert Defant and other pioneers obtained the first clear picture of the temperature, salinity, and density distributions of the deep ocean. These unprecedented observations brought about the revelation that much of the ocean is occupied by a few relatively cold and dense water masses that are formed and sink within two specific high-latitude regions: the northern North Atlantic and the Southern Ocean. Critically, as each of these water masses flows away from its formation region and pervades large areas of the globe, its initially distinct properties are eroded by mixing with surrounding waters that are, on average, lighter. As a result, the density of water masses originating at high latitudes often decreases along their path, to a point where the waters become light enough to return to the surface.

ACC 0

27.5 1000

Depth (m)

2000

28.0

3000

4000

5000

Southern Ocean 6000 − 80 − 60

North Atlantic −40

−20

0

20

40

60

Latitude (° N) Figure 1 Vertical distribution of neutral density in the Atlantic Ocean along 301 W. Contours denote isopycnal surfaces with values between 23 and 28.4 kg m 3 at intervals of 0.1 kg m 3. Colors indicate three density classes (separated by the 27.5 and 28.0 kg m 3 isopycnals) that are subject to different mixing regimes. The stratification in the abyssal ocean (purple) arises primarily from the balance between upwelling of dense water produced in the high-latitude Southern Ocean and turbulent diapycnal mixing elsewhere. In contrast, the stratification within and above the permanent pycnocline (orange) is mainly shaped by the wind- and eddy-driven subduction of near-surface water masses along isopycnals. Finally, the thick layer at the base of the permanent pycnocline, which outcrops into the upper ocean mixed layer within the Antarctic Circumpolar Current (ACC), represents a transition between the pycnocline and abyssal mixing regimes and is subject to a combination of both. The dashed arrows indicate the broad sense of the global ocean overturning driven by this set of mixing processes.


261

262

ENERGETICS OF OCEAN MIXING

This simple conceptual model lies at the heart of present views of the climatically key overturning circulation of the ocean. Implicit in the model is a competition between the removal of buoyancy from the deep ocean by dense water formation at high latitudes, and the addition of buoyancy to the deep ocean by downward mixing of light upper ocean waters, which are warmed directly by the sun. Come the second half of the century, the realization that the existence of the ocean’s overturning circulation entails the antagonistic interaction between buoyancy forcing at the sea surface and mixing in the ocean interior excited an animated debate around the driving forces of the circulation that has persisted to this day. The focus of the discussion has been on understanding the circulation’s energetics, as it is the way in which energy enters, flows through, and exits the ocean that strictly defines how the overturning circulation is driven. It is on the basis of energy considerations that early notions of surface buoyancy forcing as the governing rate-limiting process of the overturning have been challenged most convincingly in favor of ocean mixing. A central ingredient of this line of reasoning is a result put forward by Johan Sandstro¨m in 1908, stating that a fluid’s motion cannot be sustained by heating and cooling at the fluid’s surface if the source of heating lies at the same level as or above the source of cooling, such as is the case with buoyancy forcing at the sea surface. In other words, heating and cooling at a fluid’s surface do only minimal work on (i.e., input very little energy to) the fluid. Although subtle differences between the ocean and the idealized fluid that Sandstro¨m described have been argued to limit his result’s applicability to the oceanic context, it is now widely believed that mixing processes constitute the primary driving force of the overturning circulation. In this prevalent view, buoyancy forcing at the sea surface exerts an important influence on the structure of the ocean’s overturning and stratification that is particularly pronounced in transient oceanic states associated with large-scale climatic change, but cannot by itself sustain the circulation indefinitely. Thus, the problem of understanding how the overturning circulation is driven can be reduced to that of determining the energetics of ocean mixing.

The Global Ocean’s Energy Budget In order to fully appreciate the intimate link between the overturning circulation and mixing in the ocean interior, it is helpful to consider the global ocean’s energy budget. This can be formally synthesized in the global budgets of kinetic, potential, and internal

energy, which can respectively be written as: Z Z Z @=@t

rK dV ¼

Z Z Z Z

rKðu us Þ n dA ½pu þ mrK n dA

CK2P þ CI2K CK-I ½1 Z Z Z @=@t

rP dV ¼

Z Z

rPðu us Þ n dA Z Z Z þ r@Ptide =@t dV þ CK2P

½2

and Z Z Z @=@t

rI dV ¼

Z Z Z Z

rIðu us Þ n dA ½Frad rcp kT rT

@H=@S rkS rS n dA CI2K þ CK-I

½3

In these expressions, K, P, and I are the kinetic, potential, and internal energies per unit mass. The terms on the left-hand side of each equation denote the rate of change with time (t) of K, P, and I scaled by the water’s potential density (r) and integrated over the global ocean volume. These terms equal zero in the steady-state limit that is relevant to our discussion. In turn, the first terms on the equations’ right-hand sides describe the advection of the various forms of energy through the ocean surface, with u indicating the oceanic velocity, us the velocity of the free ocean surface, n a unit vector normal to that surface, and the integral being taken over the global ocean surface area. These advective terms are thought to constitute a significant source of energy to the ocean, but much of it is expended in small-scale turbulence within the upper ocean mixed layer and does not penetrate into our domain of interest, the stratified ocean interior. The second terms on the equations’ right-hand sides represent the three remaining candidate sources of the ocean interior’s energy. The term in the kinetic energy equation stands for the work done on the ocean by differential pressure (p) and viscous stresses (m is the kinematic viscosity of seawater) associated with the wind blowing on the sea surface. The differential pressure contribution is the dominant one. The term in the potential energy equation denotes the transfer of energy (expressed as a time-varying potential energy per unit mass, Ptide) from the Earth–



Moon–Sun system to the ocean by the continuous tidal displacement of the oceanic mass by gravitational forces. Finally, the term in the internal energy equation embodies surface and geothermal buoyancy forcing and amalgamates three different contributions: the radiative flux of internal energy between the near-surface ocean and overlying atmosphere/ice (Frad), and the diffusive fluxes of internal energy brought about by molecular-scale mixing of temperature (T) and salinity (S) with diffusivities kT and kS (cp and H are the specific heat capacity of seawater at constant pressure and the enthalpy of water, respectively). As advanced by Sandstro¨m’s result and reiterated by most (though not all) available recent estimates, the net buoyancy work done on the ocean by exchanges with the atmosphere is likely to be minimal. Since this has also been shown to be the case for the geothermal heating contribution, the second term on the right-hand side of the internal energy equation can be neglected, and our discussion of energy supply to the ocean will hereby focus on the two outstanding sources: the winds and the tides. The remaining terms on the right-hand sides of the three expressions above indicate the processes by which energy can be converted R R R between its various ru rP dV, characterforms. CK2P, defined as izes the transformation of kinetic energy into potential energy (or vice versa) associated with the raising or lowering of the ocean’s center of mass by advection. Although this term is often important in regional energy budgets, it averages out to a negRligible R R value in the global budget. CI2K is defined as p r u dV and represents a bidirectional transfer between the internal and kinetic energy pools due to the compressibility of seawater, which causes density to vary with pressure. This term has been estimated to be small away from the upper ocean mixed layer. energy conversion term CK-I is the only irreversible RRR and is defined as re dV, where e is the rate at which internal energy (heat) is produced by the viscous dissipation of turbulent kinetic energy per unit mass. This process represents the only significant sink of kinetic (and, indirectly, potential) energy in the ocean, and must therefore balance the energy input by winds and tides. Thus, the dominant global energy budget for the ocean interior can be synthesized as Z Z

ru n dA Z Z Z þ r@Ptitde =@t dVECK-I

½4

The validity of this balance in the characterization of

263

the ocean’s kinetic and potential energy sources and sinks is widely accepted by oceanographers. Nonetheless, establishing the physical controls and sensitivities of the overturning circulation demands that the flow of energy through the ocean be understood as well. It is the physical means of this energy flow that has been the focus of the ocean mixing debate in recent decades. In the following, we review the two most salient views of the subject to date, and provide an outlook on the major avenues of future development.

The Traditional Paradigm of Ocean Mixing: The Abyssal Ocean The longest-standing and most influential paradigm of ocean mixing and its driving of the overturning circulation was first formulated by Walter Munk in 1966. The paradigm describes how the ocean stratification below a nominal depth of 1000 m (i.e., below the ocean’s permanent pycnocline) may be explained by a simple one-dimensional balance between the upwelling (at a rate of c. 1 10 7 m s 1 or 3 m yr 1) of dense abyssal waters formed at high latitudes, and the downward turbulent mixing (at a rate defined by a turbulent diffusivity kr of c. 1 10 4 m2 s 1) of lighter overlying waters. In energetic terms, the balance is established between a decrease in the ocean’s potential energy associated with high-latitude production of dense waters, which lowers the ocean’s center of mass, and a compensating potential energy increase brought about by the lightening of those waters by turbulent diapycnal (i.e., across density surfaces) mixing as they upwell, which restores the ocean’s center of mass to its original level. A key fact that is made evident in this view is that, when oceanic turbulence ensues, not all the turbulent kinetic energy is dissipated into internal energy, but a fraction of it is expended in mixing water masses of different densities and thus leads to a vertical buoyancy flux. The relationship between the turbulent diapycnal diffusivity kr and the rate of turbulent kinetic energy dissipation e may then be expressed as kp ¼ GeN 2

½5

where the buoyancy frequency N ¼ (gr 1 @r/@z)1/2 is a measure of the stratification, g is the acceleration due to gravity, and G is the so-called mixing efficiency, commonly (and somewhat controversially) thought to be about 0.2. Using [5], it has been shown that driving the global overturning circulation across the observed ocean stratification requires that 2–3 TW is dissipated


264


by turbulence in the ocean interior, and that c. 0.5 TW is consumed by turbulent mixing in raising the ocean’s center of mass. The plausibility of this ocean mixing paradigm is suggested by the broad correspondence between the power required to support it (2–3 TW) and estimates of the rate at which work is done on the ocean by winds and tides. The wind contribution is thought to be very large, perhaps on the order of 10 TW, but an overwhelming fraction of this is likely dissipated within the upper ocean mixed layer or radiated as

surface waves toward the coastal boundaries where the waves’ energy is dissipated. The principal pathway for wind energy to enter the interior ocean is, in all likelihood, the wind work on the surface geostrophic flow (i.e., on the oceanic general circulation), which has been shown to occur at a rate of c. 0.8 TW and to be focused on the Antarctic Circumpolar Current (ACC), the broad, eastwardflowing current system that circumnavigates the Southern Ocean (Figure 2). Approximately 80% of the global wind work on the general circulation is

(a) 60° N

0° N

60° S 0° E

60° E

120° E

−24 −21 −18 −15 −12

−9

180° E

−6

−3

0

240° E

3

6

9

300° E

12

15

18

21

360° E

24

(b) 60° N

0° N

60° S 0° E

60° E

−8

−7

120° E

−6

−5

−4

−3

180° E

−2

−1

0

240° E

1

2

3

360° E

300° E

4

5

6

7

8

Figure 2 (a) Eastward and (b) northward components of the wind work on the oceanic general circulation in units of 10 3 W m 2, estimated using 4 years of satellite sea level measurements and atmospheric reanalysis winds. Reproduced from Wunsch C (1998) The work done by the wind on the oceanic general circulation. Journal of Physical Oceanography 28(11): 2332–2340.



thought to enter the ocean in the ACC. The manner in which this energy flows in the ocean interior and its ultimate fate represent one of the most significant unknowns in the ocean mixing problem. It is generally accepted that the bulk of the wind energy input to the general circulation is transferred to mesoscale eddies via the action of baroclinic instability (a further 0.2 TW may be transferred from the wind to the mesoscale eddy field directly, as eddies are generated by variable wind forcing). However, the energy’s subsequent pathway toward dissipation is uncertain. The traditional paradigm of ocean mixing proposes that a large fraction of the energy in the large-scale circulation and the mesoscale eddies may be eventually passed onto the ocean’s ubiquitous field of internal waves, through one or several poorly understood energy transfer processes. These include the generation of internal waves by geostrophic flow over small-scale topography; the spontaneous emission of internal waves by loss of geostrophic balance in mesoscale motions; and the nonlinear coupling between mesoscale eddies and internal waves propagating through them. Once in the internal wave field, energy is rapidly cascaded to increasingly smaller scales, to a point where wave breaking and turbulence ensue and a large fraction (1 – G) of the energy is dissipated through turbulence to heat. Aside from internal wave processes, it is also thought that a potentially large proportion of the wind work on the general circulation may be dissipated in turbulence generated by flows over sills within spatially confined abyssal passages, fracture zones, and midocean ridge canyons, although estimates of this contribution vary widely. A second significant mechanism via which the wind supplies energy to the ocean interior is the wind work on upper ocean inertial motions. As the wind blows on the sea surface, it generates upper ocean mixed layer currents that rotate at the local inertial frequency and can force downward- and equatorward-propagating near-inertial internal waves. The magnitude of the wind work on upper ocean inertial motions has been estimated as 0.5 TW, although energy losses to turbulence at the base of the mixed layer mean that this figure is likely to be an overestimate of the rate at which near-inertial internal waves transport energy into the ocean interior. Much of the wind work on upper ocean inertial motions occurs at mid-latitudes and exhibits a marked seasonal cycle (Figure 3), being primarily forced by winter storms. Together with the wind work on the general circulation, tides represent the primary source of the energy required to sustain ocean mixing. The rate at which the sun and the moon work on the ocean via

265

tidal forces has been estimated to be as large as 3.5 TW, but a substantial fraction of this energy (c. 2.6 TW) is dissipated on shallow continental shelves and does not access the ocean interior. The remaining c. 0.9 TW enters the deep ocean as a barotropic tide that forces flow over rough and steep topography and, in doing so, generates internal waves of tidal periodicity (internal tides) and boundary layer turbulence. The spatial distribution of this generation process is patchy, with enhanced barotropic tidal dissipation rates found over midocean ridges, continental slopes, and other elongated features such as island arcs (Figure 4). Although the bulk of tidally induced mixing occurs in the close vicinity of the generating topography, there is observational evidence of low-mode internal tides being able to transmit their energies over long distances and support turbulent mixing many hundreds of kilometres away from their generation site. Current estimates suggest that this process accounts for c. 0.2 TW, a small yet significant fraction of the total tidal energy input to the ocean interior. Recent observations suggest that a further noteworthy contribution to tidal energy dissipation may be associated with sill overflow turbulence within mid-ocean ridge canyons and other canyon-like topographic features. The final potentially significant source of energy to the ocean interior is also the most surprising and uncertain: the kinetic energy input by the marine biosphere. Net primary production in the euphotic zone produces roughly 60 TW of energy bound in carbohydrates, most of which is used in chemical form by organisms in the biosphere. However, it has been suggested that an amount of the order of 1 TW may be ultimately converted to biomechanical work done by animals swimming in the aphotic ocean. This estimate is subject to many uncertainties and remains exploratory. We conclude, therefore, that the traditional paradigm of ocean mixing, applicable below the permanent pycnocline, may be synthesized as a one-dimensional balance between the upward buoyancy flux associated with the upwelling of dense abyssal waters, and the downward buoyancy flux driven by internal wave breaking and nearboundary turbulence, whose primary energy sources are tides and the wind work on the general circulation. In the last two decades, the validity of this conceptual model has been disputed somewhat imprecisely on the basis of a growing body of measurements indicating that kr is often an order of magnitude smaller than the paradigm’s canonical value of 1 10 4 m2 s 1 within and above the permanent pycnocline, that is, outside the


266


60°

1992

40° JFM 20° 0° −20° −40° −60° 60°

0.1

1 5

50

103 F/W m−2

40° AMJ 20° 0° −20°

Latitude

−40° −60° 60° 40° JAS 20° 0° −20° −40° −60° 60° 40° OND

20° 0° −20° −40° −60° 0°

90°

180° Longitude

270°

Figure 3 Seasonal maps of the work done by the wind on upper ocean near-inertial motions in units of 10 3 W m 2, estimated using atmospheric reanalysis winds from 1992 and climatological hydrographic data. Each panel is a seasonal average over the months indicated at the upper left corner. Ice is indicated in white. Reproduced from Alford MH (2003) Improved global maps and 54-year history of wind work on ocean inertial motions. Geophysical Research Letters 30(8): 1424.

paradigm’s depth range of applicability. Despite its partially misguided motivation, this challenge has nonetheless stimulated the emergence of an alternative paradigm of ocean mixing and its energetics

that is consistent with observations of weak diapycnal mixing in the permanent pycnocline. The essential elements of this model are outlined in the following section.



60

60 40 20 0 20 40 60

(a)

40 20 0 20 40 60 0

50

100

25 20 15 10 60 40 20 0 20 40 60

150 5

200 0

5

250

300

10 15 20 25

(c)

0 350 mW m−3

60

(b)

267

50

100

150

200

3

2

1

100

150

200

1

0.5

0

0

250 1

300 2

3

350 mW m−3

(d)

40 20 0 20 40

0

50

100

150

0.8 0.6 0.4 0.2

200 0

250

300

60 350 0 mW m−3

0.2 0.4 0.6 0.8

50 1.5

250 0.5

300 1

350 mW m−3 1.5

Figure 4 Maps of dissipation of the four major tidal constituents in units of 10 3 W m 2, estimated using the TPXO.5 assimilation of satellite sea level measurements. (a) M2, (b) N2, (c) K1, (d) O1. Reproduced from Egbert GD and Ray RD (2003) Semi-diurnal and diurnal tidal dissipation from Topex/Poseidon altimetry. Geophysical Research Letters 30(17): 1907.

An Alternative Paradigm of Ocean Mixing: The Permanent Pycnocline The role of air–sea interaction in setting the stratification of the permanent pycnocline was first highlighted by Columbus Iselin in the first half of the twentieth century, when he noticed that the temperature–salinity relationships imprinted horizontally in the wintertime upper ocean mixed layer of the North Atlantic are reflected in the pycnocline’s vertical structure. This observation inspired the development of a family of conceptual models describing the renewal of water masses in the permanent pycnocline, and showing that a realistic stratification may be obtained in the upper kilometer of the ocean interior without any diapycnal mixing. A common ingredient of these models is the appeal to a three-way interaction between wind-forced (Ekman) vertical motion, isopycnal (i.e., along-density surfaces) stirring of water masses by mesoscale eddies, and atmospheric buoyancy forcing at the sea surface to explain how the subduction of relatively unmodified upper ocean waters into the ocean interior comes about. In energetic terms, the flow defined by the models is primarily driven by the wind work on the general circulation, which is then transferred to the mesoscale eddy field by the action of baroclinic instability. The models do not address the issues of how the subducted waters return to the surface and how the wind work is ultimately dissipated, that is, the mass and

energy budgets of the modeled circulation are not closed. The Southern Ocean arguably represents the most notorious manifestation of the above mechanism at work, and is thus at the heart of this alternative paradigm of ocean mixing. There, a range of density surfaces found at great depth over much of the global ocean, including some of the waters implicated in the traditional paradigm, are seen to outcrop into the upper ocean mixed layer of the ACC (Figure 1). This suggests that a considerable volume of water in the global ocean interior (roughly the layer between depths of 1000 and 2000 m) may not necessarily undergo turbulent mixing with lighter overlying water masses in order to return to the upper ocean, but that it may upwell along the steeply sloping isopycnals of the Southern Ocean instead. This notion finds support in recent studies of the Southern Ocean circulation and the dynamics of the ACC. The first indicate that upwelling of deep water (with original sources in the North Atlantic) to the surface does indeed occur over a substantial fraction of the ACC water column. Much of the upwelled water is returned northward as a wind-forced Ekman flow in the upper ocean mixed layer, where its properties are modified by air–sea interaction, and is subsequently subducted back into the interior at the ACC’s northern edge, from where it spreads northward and pervades vast areas of the global ocean’s permanent pycnocline. Studies of the dynamical balances of the ACC suggest, in turn, that


268


the upwelling of deep water may be largely sustained by the current’s vigorous mesoscale eddy field, in which nonlinear eddies act to drive a rectified southward flow across the time-mean geostrophic ACC streamlines. It thus becomes apparent that the Southern Ocean eddy field is pivotal to the two paradigms of ocean mixing presented here. On the one hand, it channels a large fraction (c. 0.65 TW, around 25–30%) of the net oceanic energy input toward dissipation scales, thereby contributing to sustain turbulent mixing across isopycnals in the traditional paradigm. On the other, it drives isopycnal upwelling of deep water masses that lie at the base of the permanent pycnocline in the mid- and low-latitude oceans. The extent to which these seemingly conflicting roles may be reconciled is unclear, but some light can be shed on the issue by considering the energetics of the alternative ocean mixing paradigm.

Lunisolar tides

The key assumption that this paradigm makes in proposing that mesoscale eddies may sustain isopycnal upwelling across the ACC is that their energy must be largely dissipated in viscous boundary layers at the ocean surface and floor, with minimal turbulent mixing anywhere in the ocean interior. The notion of bottom drag as a significant factor in the dissipation of the eddy field does find some support in the theory of geostrophic turbulence, which predicts that the evolution of newly generated mesoscale eddies involves a gradual vertical stretching that fluxes kinetic energy downward. Nonetheless, recent observations indicate that a significant fraction of the wind work on the ACC may instead contribute to sustain intense internal wave generation in areas of complex topography and ultimately lead to strong turbulent dissipation and mixing in the interior, much as described by the traditional paradigm.

Heating/ cooling

Winds

3.5TW 0.9

0.6

20

0.8 0

19 Surface waves/ turbulence 11EJ

Internal tides 0.1EJ 0.9 0.1

0.8

Upper ocean ? bdy. mixing

0.2

Beaches

Open ocean mixing

2.6

Boundary turbulence

1.0

Mesoscale eddies 13EJ ? ?

0.5

0.01

General circulation 20 YJ

?

0.7

pa

0.05

0.9 ?

?

Geothermal

0

0.2

19

Internal waves 1.4 EJ

Evap./ precip.

? 0.1

?

0.2

0.1

Bottom drag abyssal passages

Loss of balance-? 0.2

0.2 1.3

Shelves/ shallow seas

Maintenance of abyssal stratification by mixing

Figure 5 Schematic energy budget of the global ocean circulation, with uncertainties of at least factors of 2 and possibly as large as 10. The top row of boxes represents possible energy sources, with pa denoting atmospheric pressure loading. Shaded boxes are the principal energy reservoirs in the ocean, with energy values given in exajoules (EJ, 1018 J) and yottajoules (YJ, 1024 J). Fluxes into and out of the reservoirs are in terawatts (TW). The tidal input of 3.5 TW is the only accurate number here. The essential energetics consists of the conversion of c. 1.6 TW of wind work and c. 0.9 TW of tidal work into oceanic potential and kinetic energy through the generation of the large-scale circulation, and the ultimate viscous dissipation of that work into internal energy via internal wave breaking and near-boundary turbulence. The ellipse indicates the likely but uncertain importance of a loss of balance in the geostrophic mesoscale and other related processes in transferring eddy energy to the internal wave field. Dashed-dot lines indicate energy returned to the general circulation by turbulent mixing, and are first multiplied by the mixing efficiency G. Open ocean mixing by internal waves includes the upper ocean. Reproduced from Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281–314, & Annual Reviews.



Although the evidence available to date points to topographic internal wave generation in the ACC as a key agent in the transfer of eddy energy to the internal wave field, the other mechanisms introduced in the previous section are likely to enhance this transfer, and be more widely spread across the global ocean. On the whole, this points to the fascinating possibility that the two paradigms of ocean mixing presented here may be physically coupled, and suggests that it may no longer be appropriate to consider diapycnal and isopycnal water mass pathways in isolation. The emerging picture of ocean mixing is thus best described by the combination of two spatially and physically intertwined ‘pycnocline isopycnal’ and ‘abyssal diapycnal’ regimes.

Conclusion We conclude that the oceanic stratification and overturning circulation owe their existence to mechanical (rather than buoyancy) forcing. This principle is reflected in a state-of-the-art synthesis of the energy budget of the global ocean shown in Figure 5. The essential energetics consists of the conversion of 2–3 TW of wind and tidal work into oceanic potential and kinetic energy through the generation of the large-scale circulation, and the ultimate viscous dissipation of that work into internal energy via internal wave breaking and near-boundary turbulence. There are many uncertainties regarding the energy flow between the large-scale circulation and the small dissipation scales, but available evidence points to the existence of two physically coupled, spatially overlapping ocean mixing regimes. In the abyssal ocean, at depths in excess of c. 1000 m, the circulation is driven by turbulent mixing, which allows dense waters to upwell across the stable stratification and acts to counteract the decrease of the ocean’s potential energy brought about by high-latitude dense-water production. In contrast, the waters above and in the vicinity of the ocean’s permanent pycnocline, that is, roughly in the upper 2000 m, tend to flow along isopycnals primarily in response to the release by baroclinic instability of the potential and kinetic energy imparted by the wind on the general circulation. The likely subsequent transfer of some of this energy to the internal wave field couples the pycnocline and abyssal mixing regimes physically, and so may introduce important subtleties in the way the ocean responds to climatic changes in forcing. Significant open questions remain regarding all aspects of how energy enters the ocean, cascades to small scales, and dissipates. These are summarized in several major avenues of future development, of which the most prominent are: (1) quantitative

269

assessment of the energy budget of the upper ocean mixed layer, and of the mechanisms regulating the flow of wind energy across its base; (2) quantification of the energy sources to the internal wave field, and of the processes regulating its rate of dissipation; (3) determination of the mechanisms responsible for coupling internal waves and mesoscale eddies and for dissipating the latter; (4) evaluation of the global significance of sill overflow turbulence within confined passages and mid-ocean ridge canyons; (5) assessment of the global importance of double and differential diffusion, nonlinearities in the equation of state, and biomechanical mixing. In the light of the preceding discussion, it is probable that our attempts to understand the ocean’s state in past and future climates will be dangerously misguided until these issues are resolved.

Nomenclature A cp CI2K

CK-I

CK2P

Frad

H I kr K n N p P Ptide S t T u us V x y z

area specific heat capacity of seawater at constant pressure global rate of transfer between internal and kinetic energies due to the compressibility of seawater global rate of transfer of kinetic to internal energy due to turbulent dissipation global rate of transfer between kinetic and potential energies due to advection radiative flux of internal energy between the near-surface ocean and overlying atmosphere/ice enthalpy of water internal energy per unit mass turbulent diapycnal diffusivity kinetic energy per unit mass unit vector normal to the ocean surface buoyancy frequency pressure potential energy per unit mass tidal potential energy per units mass salinity time temperature three-dimensional velocity vector three-dimensional velocity vector of the free ocean surface volume eastward coordinate northward coordinate vertical coordinate


270


G e kS kT m r r

mixing efficiency rate of turbulent kinetic energy dissipation per unit mass molecular diffusivity of salinity molecular diffusivity of temperature kinematic viscosity of seawater potential density three-dimensional gradient operator (@/@x, @/@y, @/@z)

See also Antarctic Circumpolar Current. Bottom Water Formation. Breaking Waves and Near-Surface Turbulence. Dispersion and Diffusion in the Deep Ocean. Double-Diffusive Convection. Flows in Straits and Channels. Internal Tidal Mixing. Internal Tides. Internal Waves. Mesoscale Eddies. Ocean Circulation. Ocean Circulation: Meridional Overturning Circulation. Ocean Subduction. Overflows and Cascades. Three-Dimensional (3D) Turbulence. Tidal Energy. Turbulence in the Benthic Boundary Layer. Upper Ocean Mixing Processes. Vortical Modes. Wind- and BuoyancyForced Upper Ocean. Wind Driven Circulation.

Further Reading Alford MH (2003) Improved global maps and 54-year history of wind work on ocean inertial motions. Geophysical Research Letters 30(8): 1424. Bryden HL and Nurser AJG (2003) Effects of strait mixing on ocean stratification. Geophysical Research Letters 33: 1870--1872. Egbert GD and Ray RD (2003) Semi-diurnal and diurnal tidal dissipation from Topex/Poseidon altimetry. Geophysical Research Letters 30(17): 1907. Gnanadesikan A (1999) A simple predictive model for the structure of the oceanic pycnocline. Science 283: 2077--2079. Hughes CW (2002) Oceanography: An extra dimension to mixing. Nature 416: 136--139.

Hughes GO and Griffiths RW (2006) A simple convection model of the global overturning circulation, including effects of entrainment into sinking regions. Ocean Modelling 12: 46--79. Kunze E, Firing E, Hummon JM, Chereskin TK, and Thurnherr AM (2006) Global abyssal mixing inferred from lowered ADCP shear and CTD strain profiles. Journal of Physical Oceanography 36: 1553--1576. Marshall J and Radko T (2006) A model of the upper branch of the meridional overturning circulation of the Southern Ocean. Progress in Oceanography 70: 331--345. Munk WH and Wunsch C (1998) Abyssal recipes II: Energetics of tidal and wind mixing. Deep-Sea Research I 45: 1977--2010. Naveira Garabato AC, Stevens DP, Watson AJ, and Roether W (2007) Short-circuiting of the overturning circulation in the Antarctic Circumpolar Current. Nature 447: 194--197. Polzin KL, Toole JM, Ledwell JR, and Schmitt RW (1997) Spatial variability of turbulent mixing in the abyssal ocean. Science 276: 93--96. Rudnick DL, Boyd TJ, Brainard RE, et al. (2003) From tides to mixing along the Hawaiian Ridge. Science 301: 355--357. Samelson RM (2004) Simple mechanistic models of middepth meridional overturning. Journal of Physical Oceanography 34: 2096--2103. St. Laurent L and Simmons H (2006) Estimates of power consumed by mixing in the ocean interior. Journal of Climate 19: 4877--4889. Toggweiler JR and Samuels B (1998) On the ocean’s large-scale circulation near the limit of no vertical mixing. Journal of Physical Oceanography 28: 1832--1852. Webb DJ and Suginohara N (2001) Oceanography: Vertical mixing in the ocean. Nature 409: 37. Wunsch C (1998) The work done by the wind on the oceanic general circulation. Journal of Physical Oceanography 28(11): 2332--2340. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281--314.


EQUATORIAL WAVES A. V. Fedorov and J. N. Brown, Yale University, New Haven, CT, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction It has been long recognized that the tropical thermocline (the sharp boundary between warm and deeper cold waters) provides a wave guide for several types of large-scale ocean waves. The existence of this wave guide is due to two key factors. First, the mean ocean vertical stratification in the tropics is perhaps greater than anywhere else in the ocean (Figures 1 and 2), which facilitates wave propagation. Second, o since the Coriolis parameter vanishes exactly at 0 of latitude, the equator works as a natural boundary,

30° E

(a)

60° E

90° E

120° E

150° E

180°

suggesting an analogy between coastally trapped and equatorial waves. The most well-known examples of equatorial waves are eastward propagating Kelvin waves and westward propagating Rossby waves. These waves are usually observed as disturbances that either raise or lower the equatorial thermocline. These thermocline disturbances are mirrored by small anomalies in sea-level elevation, which offer a practical method for tracking these waves from space. For some time the theory of equatorial waves, based on the shallow-water equations, remained a theoretical curiosity and an interesting application for Hermite functions. The first direct measurements of equatorial Kelvin waves in the 1960s and 1970s served as a rough confirmation of the theory. By the 1980s, scientists came to realize that the equatorial waves, crucial in the response of the tropical ocean

150° W 120° W

90° W

60° W

30° W

0° E

30° E

10−2 N m−2

5 0 −5 −10

Zonal wind stress Indian

(b)

Pacific

Atlantic

0

Depth (m)

100

28 26 24 22 20 18 16 14

14

200

12

26 242 2

26

16

24 16

14

28 26 24 22 18 16

14

14 12

12

300

400 30° E

60° E

90° E

120° E

150° E

180°

150° W 120° W

90° W

60° W

30° W

0° E

30° E

Figure 1 The thermal structure of the upper ocean along the equator: (a) the zonal wind stress along the equator; shading indicates the standard deviation of the annual cycle. (b) Ocean temperature along the equator as a function of depth and longitude. The east– west slope of the thermocline in the Pacific and the Atlantic is maintained by the easterly winds. (b) From Kessler (2005).


271

272

EQUATORIAL WAVES

Pacific 0

.0

26

24.0 22.0 20.0 18.0 16.0 14.0

200 12 .0

Depth (m)

100

10

.0

300

400 10° S

0°

10° N

Figure 2 Ocean temperature as a function of depth and latitude in the middle of the Pacific basin (at 1401 W). The thermocline is particularly sharp in the vicinity of the equator. Note that the scaling of the horizontal axis is different from that in Figure 1. Temperature data are from Levitus S and Boyer T (1994) World Ocean Atlas 1994, Vol. 4: Temperature NOAA Atlas NESDIS4. Washington, DC: US Government Printing Office.

Temperature (°C)

28 27 26 25 24 1900

1920

1940

1960

1980

2000

Figure 3 Interannual variations in sea surface temperatures (SSTs) in the eastern equatorial Pacific shown on the background of decadal changes (in 1C). The annual cycle and high-frequency variations are removed from the data. El Ninõ conditions correspond to warmer temperatures. Note El Ninõ events of 1982 and 1997, the strongest in the instrumental record. From Fedorov AV and Philander SG (2000) Is El Ninõ changing? Science 288: 1997–2002.

to varying wind forcing, are one of the key factors in explaining ENSO – the El Ninõ-Southern Oscillation phenomenon. El Ninõ, and its complement La Ninã, have physical manifestations in the sea surface temperature (SST) of the eastern equatorial Pacific (Figure 3). These climate phenomena cause a gradual horizontal redistribution of warm surface water along the equator: strong zonal winds during La Ninã years pile up the warm water in the west, causing the thermocline to slope downward to the west and exposing cold water to the surface in the east (Figure 4(b)). During an El Ninõ, weakened zonal winds permit the warm water to flow back eastward so that the thermocline becomes more

horizontal, inducing strong warm anomalies in the SST (Figure 4(a)). The ocean adjustment associated with these changes crucially depends on the existence of equatorial waves, especially Kelvin and Rossby waves, as they can alter the depth of the tropical thermocline. This article gives a brief summary of the theory behind equatorial waves, the available observations of those waves, and their role in ENSO dynamics. For a detailed description of El Ninõ phenomenology the reader is referred to other relevant papers in this encyclopedia El Ninõ Southern Oscillation (ENSO)El Ninõ Southern Oscillation (ENSO) Modelsand Pacific Ocean Equatorial Currents.


EQUATORIAL WAVES

273

(a) 140° E 160° E

180°

160° W 140° W 120° W 100° W

0

32 28

28

Depth (m)

100

24

22 20

20

18 16

14

200

16

12

12 8

300

4 0

400 (b) 140° E 160° E

180°

160° W 140° W 120° W 100° W

32

0

28 Depth (m)

100

24

28

20 200

16

14

12

12

8

300

4 400

10

0

Figure 4 Temperatures (1C) as a function of depth along the equator at the peaks of (a) El Ninõ (Jan. 1998) and (b) La Ninã (Dec. 1999). From the TAO data; see McPhaden MJ (1999) Genesis and evolution of the 1997–98 El Ninõ. Science 283: 950–954.

It is significant that ENSO is characterized by a spectral peak at the period of 3–5 years. The timescales associated with the low-order (and most important dynamically) equatorial waves are much shorter than this period. For instance, it takes only 2–3 months for a Kelvin wave to cross the Pacific basin, and less than 8 months for a first-mode Rossby wave. Because of such scale separation, the properties of the ocean response to wind perturbations strongly depend on the character of the imposed forcing. It is, therefore, necessary to distinguish the following. 1. Free equatorial waves which arise as solutions of unforced equations of motion (e.g., free Kelvin and Rossby waves), 2. Equatorial waves forced by brief wind perturbations (of the order of a few weeks). In effect, these waves become free waves as soon as the wind perturbation has ended, 3. Equatorial wave-like anomalies forced by slowly varying periodic or quasi-periodic winds reflecting ocean adjustment on interannual timescales. Even though these anomalies can be represented mathematically as a superposition of continuously forced Kelvin and Rossby waves of different modes, the properties of the superposition (such

as the propagation speed) can be rather different from the properties of free waves.

The Shallow-Water Equations Equatorial wave dynamics are easily understood from simple models based on the 1 12-layer shallowwater equations. This approximation assumes that a shallow layer of relatively warm (and less dense) water overlies a much deeper layer of cold water. The two layers are separated by a sharp thermocline (Figure 5), and it is assumed that there is no motion in the deep layer. The idea is to approximate the thermal (and density) structure of the ocean displayed in Figures 1 and 2 in the simplest form possible. The momentum and continuity equations, usually referred to as the reduced-gravity shallow-water equations on the b-plane, are ut þ g0 hx byv ¼ tx =rD as u

½1

vt þ g0 hy þ byu ¼ ty =rD as v

½2

ht þ Hðux þ vy Þ ¼ as h

½3


274

EQUATORIAL WAVES

H1=H

h(x,y,t)

+Δ

H2

z

y x

Figure 5 A sketch of the 1 12-layer shallow-water system with the rigid-lid approximation. H1/H2{1. The mean depth of the thermocline is H. The x-axis is directed to the east along the equator. The y-axis is directed toward the North Pole. The mean east–west thermocline slope along the equator is neglected.

These equations have been linearized, and perturbations with respect to the mean state are considered. Variations in the mean east–west slope of the thermocline and mean zonal currents are neglected. The notations are conventional (some are shown in Figure 5), with u, v denoting the ocean zonal and meridional currents, H the mean depth of the thermocline, h thermocline depth anomalies, tx and ty the zonal and meridional components of the wind stress, r mean water density, Dr the difference between the density of the upper (warm) layer and the deep lower layer, g0 ¼ gDr/r the reduced gravity. D is the nominal depth characterizing the effect of wind on the thermocline (frequently it is assumed that D ¼ H). The subscripts t, x, and y indicate the respective derivatives. The system includes simple Rayleigh friction in the momentum equations and a simple linear parametrization of water entrainment at the base of the mixed layer in the continuity equation (terms proportional to as). Some typical values for the equatorial Pacific are Dr/r ¼ 0.006; H ¼ 120 m (see Figure 1); D ¼ 80 m. The rigid-lid approximation is assumed (i.e., to a first approximation the ocean surface is flat). However, after computing h, one can calculate small changes in the implied elevation of the free surface as Z ¼ hDr/r. It is this connection that allows us to estimate changes in the thermocline depth from satellite measurements by measuring sealevel height. The boundary conditions for the equations are no zonal flow (u ¼ 0) at the eastern and western boundaries and vanishing meridional flow far away from the equator. The former boundary conditions are sometimes modified by decomposing u into different wave components and introducing reflection coefficients (smaller than unity) to account for a partial reflection of waves from the boundaries.

It is apparent that the 1 12-layer approach leaves in the system only the first baroclinic mode and eliminates barotropic motion (the first baroclinic mode describes a flow that with different velocities in two layers, barotropic flow does not depend on the vertical coordinate). This approximation filters out higher-order baroclinic modes with more elaborate vertical structure. For instance, the equatorial undercurrent (EUC) is absent in this model. Observations and numerical calculations show that to represent the full vertical structure of the currents and ocean response to winds correctly, both the firstand second-baroclinic modes are necessary. Nevertheless, the shallow-water equations within the 1 12layer approximation remain very successful and are used broadly in the famous Cane-Zebiak model of ENSO and its numerous modifications. For many applications, the shallow-water equations are further simplified to filter out short waves: in the long-wave approximation the second momentum equation (eqn [2]) is replaced with a simple geostrophic balance: g0 hy þ byu ¼ 0

½4

The boundary condition at the western boundary is then replaced with the no-net flow requirement R udy ¼ 0 . It is noteworthy that the shallow-water equations can also approximate the mean state of the tropical ocean (if used as the full equations for mean variables, rather than perturbations from the mean state). In that case the main dynamic balance along the equator is that between the mean trade winds and the mean (climatological) slope of the thermocline (damping neglected) g0 hx B tx =rD

½5

This balance implies the east–west difference in the thermocline depth along the equator of about 130 m in the Pacific consistent with Figure 1.

Free-Wave Solutions of the ShallowWater Equations First, we consider the shallow-water eqns [1]–[3] with no forcing and no dissipation. The equations have an infinite set of equatorially trapped solutions (with v-0 for y-7N). These are free equatorial waves that propagate back and forth along the equator.


EQUATORIAL WAVES

275

as in [9] we obtain a single equation for v˜(y):

Kelvin Waves

Kelvin waves are a special case when the meridional velocity vanishes everywhere identically (v ¼ 0) and eqns [1]–[3] reduce to ut þ g0 hx ¼ 0

½6

g0 hy þ byu ¼ 0

½7

ht þ Hux ¼ 0

½8

Looking for wave solutions of [6]–[8] in the form iðkxotÞ ˜ ½u; v; h ¼ ½uðyÞ; vðyÞ; hðyÞe ˜ ˜

vðyÞ ¼ 0; ˜

o>0

½9 ½10

we obtain the dispersion relation for frequency o and wave number k o2 ¼ g0 Hk2

½11

and a first-order ordinary differential equation for the meridional structure of h dh˜ bk ¼ yh˜ dy o

½12

The only solution of [11] and [12] decaying for large y, called the Kelvin wave solution, is h ¼ h0 e

ðb=2cÞy2 iðkxotÞ

e

0

½13

1/2

where the phase speed c ¼ (g H) and o ¼ ck (h0 is an arbitrary amplitude). Thus, Kelvin waves are eastward propagating (o/k40) and nondispersive. The second solution of [11] and [12], the one that propagates westward, would grow exponentially for large y and as such is disregarded. Calculating the Kelvin wave phase speed from typical parameters used in the shallow-water model gives c ¼ 2.7 m s 1 which agrees well with the measurements. The meridional scale with which these solutions decay away from the equator is the equatorial Rossby radius of deformation defined as LR ¼ ðc=bÞ1=2

½14

which is approximately 350 km in the Pacific Ocean, so that at 51 N or 51 S the wave amplitude reduces to 30% of that at the equator. Rossby, Poincare, and Yanai Waves

Now let us look for the solutions that have nonzero meridional velocity v. Using the same representation

! o2 bk b2 2 2 k y v˜ ¼ 0 c2 o c2

d2 v˜ þ dy2

½15

The solutions of [15] that decay far away from the equator exist only when an important constraint connecting its coefficients is satisfied: o2 bk b ¼ ð2n þ 1Þ k2 2 c o c

½16

where n ¼ 0, 1, 2, 3, y. This constraint serves as a dispersion relation o ¼ o(k,n) for several different types of equatorial waves (see Figure 6), which include 1. 2. 3. 4.

Gravity-inertial or Poincare waves n ¼ 1, 2, 3, y Rossby waves n ¼ 1, 2, 3, y Rossby-gravity or Yanai wave n ¼ 0 Kelvin wave n ¼ 1.

These waves constitute a complete set and any solution of the unforced problem can be represented as a sum of those waves (note that the Kelvin wave is formally a solution of [15] and [16] with v ¼ 0, n ¼ 1). Let us consider several important limits that will elucidate some properties of these waves. For high frequencies we can neglect bk/o in [16] to obtain o2 ¼ c2 k2 þ ð2n þ 1Þbc

½17

where n ¼ 1, 2, 3, y. This is a dispersion relation for gravity-inertial waves, also called equatorially trapped Poincare waves. They propagate in either direction and are similar to gravity-inertial waves in mid-latitudes. For low frequencies we can neglect o2/c2 in [16] to obtain bk o¼ 2 k þ ð2n þ 1Þb=c

½18

with n ¼ 1, 2, 3, y. These are Rossby waves similar to their counterparts in mid-latitudes that critically depend on the b-effect. Their phase velocity (o/k) is always westward (o/ko0), but their group velocity @o/@k can become eastward for high wave numbers. The case n ¼ 0 is a special case corresponding to the so-called mixed Rossby-gravity or Yanai wave. Careful consideration shows that when the phase velocity of those waves is eastward (o/k40), they behave like gravity-inertial waves and [17] is satisfied, but when the phase velocity is westward (o/ko0), they behave like Rossby waves and expression [18] becomes more appropriate.


276

EQUATORIAL WAVES

/(2c)1/2 1 2

Poincare waves

21 0

3

2

Kelvin wave 1 0

Rossby waves

1

k (c /2)1/2

2 −3

−2

−1

0

1

2

3

Figure 6 The dispersion relation for free equatorial waves. The axes show nondimensionalized wave number and frequency. The blue box indicates the long-wave regime. n ¼ 0 indicates the Rossby-gravity (Yanai) wave. Kelvin wave formally corresponds to n ¼ 1. From Gill AE (1982) Atmosphere-Ocean Dynamics, 664pp. New York: Academic Press.

The meridional structure of the solutions of eqn [15] corresponding to the dispersion relation in [16] is described by Hermite functions: pffiffiffi n 1=2 2 Hn ðb=cÞ1=2 y eðb=2cÞy ½19 v˜ ¼ p2 n! where Hn(Y) are Hermite polynomials (n ¼ 0, 1, 2, 3, y), H0 ¼ 1; H1 ¼ 2Y; H2 ¼ 4Y 2 2; H3 ¼ 8Y 3 12Y; H4 ¼ 16Y 4 48Y 2 þ 12; y; ½20

It is Rossby waves of low odd numbers and Kelvin waves that usually dominate the solutions for large-scale tropical problems. This suggests that solving the equations can be greatly simplified by filtering out Poincare and short Rossby waves. Indeed, the long-wave approximation described earlier does exactly that. Such an approximation is equivalent to keeping only the waves that fall into the small box in Figure 6, as well as a remnant of the Yanai wave, and then linearizing the dispersion relations for small k. This makes long Rossby wave nondispersive, each mode satisfying a simple dispersion relation with a fixed n:

and 1=2

Y ¼ ðb=cÞ

y

½21

These functions as defined in [19]–[21] are orthonormal. The structure of Hermite functions, and hence of the meridional flow corresponding to different types of waves, is plotted in Figure 7. Hermite functions of odd numbers (n ¼ 1, 3, 5, y) are characterized by zero meridional flow at the equator. It can be shown that they create symmetric thermocline depth anomalies with respect to the equator (e.g., a firstorder Rossby wave with n ¼ 1 has two equal maxima in the thermocline displacement on each side of the equator). Hermite functions of even numbers generate cross-equatorial flow and create thermocline displacement asymmetric with respect to the equator (i.e., with a maximum in the thermocline displacement on one side of the equator, and a minimum on the other).

o¼

c k; 2n þ 1

n ¼ 1; 2; 3; y

½22

Consequently, the phase speed of Rossby waves of different modes is c/3, c/5, c/7, etc. The phase speed of the first Rossby mode with n ¼ 1 is c/3, that is, one-third of the Kelvin wave speed. It takes a Kelvin wave approximately 2.5 months and a Rossby wave 7.5 months to cross the Pacific. The higher-order Rossby modes are much slower. The role of Kelvin and Rossby waves in ocean adjustment is described in the following sections.

Ocean Response to Brief Wind Perturbations First, we will discuss the classical problem of ocean response to a brief relaxation of the easterly trade winds. These winds normally maintain a strong east–west thermocline slope along the equator


EQUATORIAL WAVES

0.8

277

0.8 1 3 5

0.6 0.4

0.4

0.2

0.2

0

0

−0.2

−0.2

−0.4

−0.4

−0.6

−0.6 −4

−2

0

2

0

0.6

−4

4

−2

0

y/LR

2

4

2

4

y/LR

Figure 7 The meridional structure of Hermite functions corresponding to different equatorial modes (except for the Kelvin mode). The meridional velocity v is proportional to these functions. Left: Hermite functions of odd numbers (n ¼ 1, 3, 5, y) with no meridional flow crossing the equator. The flow is either converging or diverging away from the equator, which produces a symmetric structure (with respect to the equator) of the thermocline anomalies. Right: Hermite functions of even numbers (n ¼ 0, 2, 4, y) with nonzero cross-equatorial flow producing asymmetric thermocline anomalies.

20° N

0

Latitude

10° N 0.2

EQ 10° S

Time (years)

0.4 20° S 140° E 0.6

180° E

140° W

100° W

Longitude 20° N

0.8 Latitude

10° N 1

EQ 10° S

1.2

20° S 140° E

180° E

140° W

100° W

140° E

−10

0

10

140° W

100° W

Longitude

Longitude −20

180° E

20

30

−20

−10

0

10

20

30

Figure 8 Ocean response to a brief westerly wind burst (WWB) occurring around time t ¼ 0 in a shallow-water model of the Pacific. Left: a Hovmoller diagram of the thermocline depth anomalies along the equator (in meters). Note the propagation and reflection of Kelvin and Rossby waves (the signature of Rossby waves on the equator is usually weak and rarely seen in the observations). Right: the spatial structure of the anomalies at times indicated by the white dashed lines on the left-side panel. The arrows indicate the direction of wave propagation. The wind-stress perturbation is given by t ¼ twwb exp[ (t/t0)2 (x/Lx)2 (y/Ly)2]. Red corresponds to a deeper thermocline, blue to a shallower thermocline. The black dashed ellipse indicates the timing and longitudinal extent of the WWB.

and their changes therefore affect the ocean state. Westerly wind bursts (WWBs) that occur over the western tropical Pacific in the neighborhood of the dateline, lasting for a few weeks to a

month, are examples of such occurrences. (Early theories treated El Ninõ as a simple response to a wind relaxation caused by a WWB. Arguably, WWBs may have contributed to the development of


278

EQUATORIAL WAVES

El Ninõ in 1997, but similar wind events on other occasions failed to have such an effect.) As compared to the timescales of ocean dynamics, these wind bursts are relatively short, so that ocean adjustment occurs largely when the burst has already ended. In general, the wind bursts have several effects on the ocean, including thermodynamic effects modifying heat and evaporation fluxes in the western tropical Pacific. The focus of this article is on the dynamic effects of the generation, propagation, and then boundary reflection of Kelvin and Rossby waves. The results of calculations with a shallow-water model in the long-wave approximation are presented next, in which a WWB lasting B3 weeks is applied at time t ¼ 0 in the Pacific. The temporal and spatial structure of the burst is given by 2

t ¼ twwb eðt=t0 Þ

ðy=Ly Þ2 ðxx0 Þ2 =L2x

½23

which is roughly consistent with the observations. The burst is centered at x0 ¼ 1801 W; and Lx ¼ 101; Ly ¼ 101; t0 ¼ 7 days; twwb ¼ 0.02 N m 2.

The WWB excites a downwelling Kelvin wave and an upwelling Rossby wave seen in the anomalies of the thermocline depth. Figure 8 shows a Hovmoller diagram and the spatial structure of these anomalies at two particular instances. The waves propagate with constant speeds, although in reality Kelvin waves should slow down in the eastern part of the basin where the thermocline shoals (since c ¼ (g0 H)1/2). The smaller slope of the Kelvin wave path on the Hovmoller diagram corresponds to its higher phase speed, as compared to Rossby waves. The spatial structure of the thermocline anomalies at two instances is shown on the right panel of Figure 8. The butterfly shape of the Rossby wave (meridionally symmetric, but not zonally) is due to the generation of slower, high-order Rossby waves that trail behind (higher-order Hermite functions extend farther away from the equator, Figure 7). The waves reflect from the western and eastern boundaries (in the model the reflection coefficients were set at 0.9). When the initial upwelling Rossby wave reaches the western boundary, it reflects as an equatorial upwelling Kelvin wave. When the downwelling Kelvin wave reaches the eastern boundary, a

20° N

1

Latitude

10° N 2

EQ 10° S

Time (years)

3 20° S 180° E 140° W

140° E 4

100° W

Longitude 20° N

5 Latitude

10° N 6

EQ 10° S

7

20° S 140° E

180° E

140° W

100° W

140° E

180° E

Longitude −50

0

140° W

100° W

Longitude 50

−50

0

50

Figure 9 Ocean response to oscillatory winds in a shallow-water model. Left: a Hovmoller diagram of the thermocline depth anomalies along the equator. Note the different temporal scale as compared to Figure 8. Right: the spatial structure of the anomalies at times indicated by the dashed lines on the left-side panel. Red corresponds to a deeper thermocline, blue to a shallower thermocline. The wind-stress anomaly is calculated as t ¼ t0 sin(2pt/P)*exp [ (x/Lx)2 (y/Ly)2]; P ¼ 5 years.


EQUATORIAL WAVES

number of things occur. Part of the wave is reflected back along the equator as an equatorial downwelling Rossby wave. The remaining part travels north and south as coastal Kelvin waves, apparent in the lower right panel of Figure 8, which propagate along the west coast of the Americas away from the Tropics.

The results of calculations with a shallow-water model in which a periodic sinusoidal forcing with the period P is imposed over the ocean are presented in Figure 9. The spatial and temporal structure of the forcing is given by 2

t ¼ t0 sinð2pt=PÞeðy=Ly Þ

Ocean Response to Slowly Varying Winds Ocean response to slowly varying periodic or quasiperiodic winds is quite different. The relevant zonal dynamical balance (with damping neglected) is ðut byvÞ þ g0 hx ¼ tx =rD

½24

It is the balance between the east–west thermocline slope and the wind stress that dominates the equatorial strip. Off the equator, however, the local wind stress is not in balance with the thermocline slope as the Coriolis acceleration also becomes important. (a)

279

ðxx0 Þ2 =L2x

½25

where we choose x0 ¼ 1801 W; Lx ¼ 401; Ly ¼ 101; and t0 ¼ 0.02 N m 2; and P ¼ 5 years. This roughly approximates to interannual wind stress anomalies associated with ENSO. Figure 9 shows a Hovmoller diagram and the spatial structure of the ocean response at two particular instances. The thermocline response reveals slow forced anomalies propagating eastward along the equator. As discussed before, mathematically they can be obtained from a supposition of Kelvin and Rossby modes; however, the individual free waves are implicit and cannot be identified in the response. At the peaks of the anomalies, the spatial structure of the ocean response is characterized by

North

Shallow Wind Stress

EQ

Deep thermocline Warm SST Shallow

South (b)

North

Deep Wind EQ

Shallow thermocline Cold SST

Stress Deep

South Figure 10 A schematic diagram that shows the spatial (longitude–latitude) structure of the coupled ‘delayed oscillator’ mode. Arrows indicate anomalous wind stresses, colored areas changes in thermocline depth. The sketch shows conditions during (a) El Ninõ and (b) La Ninã. The off-equatorial anomalies are part of the ocean response to varying winds (cf. Figure 9). While the spatial structure of the mode resembles a pair of free Kelvin and Rossby waves, it is not so. The transition from (a) to (b) includes the shallow offequatorial anomalies in thermocline depth slowly feeding back to the equator along the western boundary and then traveling eastward to reemerge in the eastern equatorial Pacific and to push the thermocline back to the surface. It may take, however, up to several years, instead of a few months, to move from (a) to (b). From Fedorov AV and and Philander SG (2001) A stability analysis of tropical ocean–atmosphere interactions: Bridging measurements and theory for El Ninõ. Journal of Climate 14(14): 3086–3101.


280

EQUATORIAL WAVES

thermocline depression or elevation in the eastern equatorial Pacific (in a direct response to the winds) and off-equatorial anomalies of the opposite sign in the western equatorial Pacific.

Conceptual Models of ENSO Based on Ocean Dynamics So far the equatorial processes have been considered strictly from the point of view of the ocean. In

particular, we have shown that wind variations are able to excite different types of anomalies propagating on the thermocline – from free Kelvin and Rossby waves generated by episodic wind bursts to gradual changes induced by slowly varying winds. However, in the Tropics variations in the thermocline depth can affect SSTs and hence the winds, which gives rise to tropical ocean–atmosphere interactions. The strength of the easterly trade winds (that maintain the thermocline slope in Figure 1) is roughly proportional to the east–west temperature

Period (years)

(a) 3

9 2.5 8

b (1/year)

2 7 1.5 6 1 5 0.5 4 0.5

1 a (1/year)

1.5

2

Growth rates (1/year)

(b) 3

0.8 2.5

0.6 0.4

b (1/year)

2

0.2 1.5

0 −0.2

1

−0.4 0.5

−0.6 0.5

1 a (1/year)

1.5

2

Figure 11 The period and the e-folding growth (decay) rates of the ENSO-like oscillation given by the delayed oscillator model dT/dt ¼ aT bT(t D) as a function of a and b; for the delay D ¼ 12 months. The solutions of the model are searched for as est, where s is a complex frequency. In the white area of the plot there are no oscillatory, but only exponentially growing or decaying solutions. At the border between the white and color areas the oscillation period goes to infinity, that is, imag(s) ¼ 0. The dashed line in (b) indicates neutral stability. Note that the period of the oscillation can be much longer than the delay D used in the equation.


EQUATORIAL WAVES

gradient along the equator. This implies a circular dependence: for instance, weaker easterly winds, during El Ninõ, result in a deeper thermocline in the eastern equatorial Pacific, weaker zonal SST gradient, and weaker winds. This is a strong positive feedback usually referred to as the Bjerknes feedback. On the other hand, the gradual oceanic response to changes in the winds (often referred to as ‘ocean memory’) provides a negative feedback and a potential mechanism for oscillatory behavior in the system. In fact, the ability of the ocean to undergo slow adjustment delayed with respect to wind variations and the Bjerknes feedback serve as a basis for one of the first conceptual models of ENSO – the delayed oscillator model. Delayed Oscillator

Zonal wind fluctuations associated with ENSO occur mainly in the western equatorial Pacific and give rise

(a)

a

to basin-wide vertical movements of the thermocline that affect SSTs mainly in the eastern equatorial Pacific. During El Ninõ, the thermocline in the east deepens resulting in the warming of surface waters. At the same time, the thermocline in the west shoals; the shoaling is most pronounced off the equator. In this coupled mode, shown schematically in Figure 10, the response of the zonal winds to changes in SST is, for practical purposes, instantaneous, and this gives us the positive Bjerknes feedback described above. Ocean adjustment to changes in the winds, on the other hand, is delayed. The thermocline anomalies off the equator slowly feed back to the equator along the western boundary and then travel eastward, reemerging in the eastern equatorial Pacific, pushing the thermocline back to the surface, and cooling the SST. It may take up to a year or two for this to occur. This mode can therefore be called as a ‘delayed oscillator’ mode.

(b)

Sverdrup transport

a~0

SSTa(+)

SSTa~0 EQ

Depth anomaly

Depth anomaly

EQ Warm water

Cold water

(c) a

(d) a~0

SSTa(−)

Sverdrup transport

SSTa~0 EQ

Depth anomaly

EQ Depth anomaly

281

Figure 12 (a–d) A sketch showing the recharge-discharge mechanism of ENSO. All quantities are anomalies relative to the climatological mean. Depth anomaly is relative to the time mean state along the equator. Dashed line indicates zero anomaly; shallow anomalies are above the dashed line and deep anomalies are below the dashed line. Thin arrows and symbol ta represent the anomalous zonal wind stress; bold thick arrows represent the corresponding anomalous Sverdrup transports. SSTa is the sea surface temperature anomaly. Oscillation progresses from (a) to (b), (c), and (d) clockwise around the panels following the roman numerals; panel (a) represents El Ninõ conditions, panel (c) indicates La Ninã conditions. Note similarities with Figure 10. Modified from Jin FF (1997) An equatorial ocean recharge paradigm for ENSO. 1. Conceptual model. Journal of the Atmospheric Sciences 54: 811–829 and Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Ninõ and La Ninã. Journal of Climate 13: 3551–3559.


282

EQUATORIAL WAVES

Tt ¼ aT þ bTðt DÞ

½26

where T is temperature, a and b are constants, t is time, and D is a constant time lag. The first term on the right-hand side of the equation represents the positive feedbacks between the ocean and atmosphere (including the Bjerknes feedback). It is the presence of the second term that describes the delayed response of the ocean that permits oscillations (the physical meaning of the delay D is the time needed for an off-equatorial anomaly in the western Pacific to converge to the equator and then travel to the eastern Pacific). The period of the simulated oscillation depends on the values of a, b, and D. Solutions of eqn [26] proportional to est give a transcendental algebraic equation for the complex frequency s s ¼ a best

½27

s ¼ sr þ isi

½28

where

The solutions of eqn [27] are shown in Figure 11 for D ¼ 1 year and different combinations of a and b.

Even though the term ‘delayed oscillator’ appears frequently in the literature, there is some confusion concerning the roles of Kelvin and Rossby waves, which some people seem to regard as the salient features of the delayed oscillator. The individual waves are explicitly evident when the winds change abruptly (Figure 6), but those waves are implicit when gradually varying winds excite a host of waves, all superimposed (Figure 7). For the purpose of deriving the delayed-oscillator equation, for instance, observations of explicit Kelvin (and for that matter individual Rossby) waves are irrelevant. The gradual eastward movement of warm water in Figure 7 (left panel) is the forced response of the ocean and cannot be a wave that satisfies the unforced equations of motion. Recharge Oscillator

The delayed oscillator gave rise to many other conceptual models based on one or another type of the delayed-action equation (the Western Pacific, Advection, Unified oscillators, just to name a few, each emphasizing particular mechanisms involved in ENSO). A somewhat different approach was used by Jin in 1997 who took advantage of the fact that free Kelvin waves cross the Pacific very quickly, which allowed him to eliminate Kelvin waves from

Eq

4° N 1000

Oct. Jul.

1995

An equation that captures the essence of this mode is

Apr. Jan.

800

Jul.

Day

600

1994

Oct.

Apr. Jan. 400 Jul.

200

1993

Oct. X

Apr. Jan.

0 120°

150° E

180°

150° W

120° −6

−4

90°

120°

−2

0

150° E 2

4

180°

150° W

120°

90°

6

Sea level (cm) Figure 13 Observations of Rossby (left) and Kelvin (right) waves. Time-longitude sections of filtered sea level variations in the Pacific Ocean along 41 N and along the equator are shown. A section along 41 S would be similar to the 41 N section. The symbols (x, triangle, and circles) correspond to the times and locations of the matching symbols in Figure 14. Note that time runs from bottom to top. Obtained from TOPEX/POSEIDON satellite data; from Chelton DB and Schlax MG (1996) Global observations of oceanic Rossby Waves. Science 272(5259): 234–238.


EQUATORIAL WAVES

consideration and derive the recharge oscillator model. The recharge oscillator theory is now one of the commonly used paradigms for ENSO. It relies on a phase lag between the zonally averaged thermocline depth anomaly and changes in the eastern Pacific SST. Consider first a ‘recharged’ ocean state (Figure 12(d)) with a deeper than normal thermocline across the tropical Pacific. Such a state is conducive to the development of El Ninõ as the deep thermocline inhibits the upwelling of cold water in the east. As El Ninõ develops (Figure 12(a)), the reduced zonal trade winds lead to an anomalous Sverdrup transport out of the equatorial region. The ocean response involves a superposition of many equatorial waves resulting in a shallower than normal equatorial thermocline and the termination of El Ninõ (Figure 12(b)). The state with a shallower mean thermocline (the ‘discharged’ state) is usually followed by a La Ninã

283

event (Figure 12(c)). During and after La Ninã the enhanced trade winds generate an equatorward flow, deepening the equatorial thermocline and eventually ‘recharging’ the ocean (Figure 12(d)). This completes the cycle and makes the ocean ready for the next El Ninõ event.

Observations of Kelvin and Rossby Waves, and El Ninõ Thirty years ago very little was known about tropical processes, but today an impressive array of instruments, the TAO array, now monitors the equatorial Pacific continuously. It is now possible to follow, as they happen, the major changes in the circulation of the tropical Pacific Ocean that accompany the alternate warming and cooling of the surface waters of the eastern equatorial Pacific associated with El Ninõ and La Ninã. Satellite-borne radiometers and

Cycle 21 (13 April 1993) 50° N 40 30 20 10 0 10 20 30 40 50° S 60° E

120°

180°

120°

60° W

0°

60° W

0°

Cycle 32 (31 July 1993) 50° N 40 30 20 10 0 10 20 30 40 50° S

60° E

120° −4

180°

120°

−2 0 2 Sea level (cm)

4

Figure 14 Observations of Rossby waves: global maps of filtered sea level variations on 13 April 1993 and 3 12 months later on 31 July. White lines indicate the wave trough. The time evolution of the equatorial Kelvin wave trough (x), the Rossby wave crest (open triangle and open circle), and the Rossby wave trough (solid circle) can be traced from the matching symbols in Figure 13. Obtained from TOPEX/POSEIDON satellite data, from Chelton and Schlax (1996).


EQUATORIAL WAVES

0

0

0

20 20

20

40 0

20

60

20

40

20

0

Time (month)

40

80

40

0

20 40 40

40 0 140° E

160° E

180° −80

−40

160° W 140° W 0

40

80

120° W 100° W

0

S O N D J F M A M J J A S O N D J F

0

20

0

0 40

60

0

higher-order waves travel much slower. The waves are forced by high-frequency wind perturbations, even though it seems likely that annual changes in the zonal winds may have also contributed to the forcing of Rossby waves. As mentioned above, there is a clear distinction between free Kelvin waves and slow wave-like anomalies associated with ENSO. This is further emphasized by the measurements in Figure 15 that contain evidence of freely propagating Kelvin waves (dashed lines in the left panel) but clearly show them to be separate from the far more gradual eastward movement of warm water associated with the onset of El Ninõ of 1997 (a dashed line in the right panel). This slow movement of warm water is the forced response of the ocean and clearly not a wave that could satisfy the unforced equations of motion. The characteristic timescale of ENSO cycle, several years, is so long that low-pass filtering is required to isolate its structure. That filtering eliminates individual Kelvin waves in the right-side panel of the figure. Whether the high- and low-frequency components of

20

20

0

40

M A M J J A S O N D

40

20

40

altimeters measure ocean temperature and sea level height almost in real time, providing information on slow (interannual) changes in the ocean thermal structure as well as frequent glimpses of swift wave propagation. Figures 13 and 14 show the propagation of fast Kelvin and Rossby waves in the Pacific as seen in the satellite altimeter measurements of the sea level height. The speed of propagation of Kelvin waves agrees relatively well with the predictions from the theory (B2.7 m s 1). The speed of the first-mode Rossby waves, however, is estimated from the observation to be 0.5–0.6 m s 1, which is somewhat lower than expected, that is, 0.9 m s 1. This appears to be in part due to the influence of the mean zonal currents. Estimated variations of the thermocline depth associated with the small changes in the sea level in Figure 13 are in the range of 75 m (stronger Kelvin waves may lead to variation up to 720 m). Note that the observed ‘Rossby wave’ in Figure 14 is actually a composition of Rossby waves of different orders;

20

0

284

140° E

160° E −80

180° −40

160° W 140° W 120° W 0

40

M J J A S O N D J F M A M J J A S O N D J F M A M J J A S O N D J F M A

100° W

80

Figure 15 Observations of Kelvin waves and El Ninõ from the TAO array. Anomalies with respect to the long-term average of the depth of the 20 1C-degree isotherm are shown before and after the development of El Ninõ of 1997. Left: 5-day averages; the time axis starts in September 1996. Right: monthly averages; the time axis starts in May 1996 (cf. Figures 8 and 9). The dashed lines in the leftside panel correspond to Kelvin waves excited by brief WWBs and rapidly traveling across the Pacific. The dashed lines in the rightside panel show the slow eastward progression of warm and cold temperature anomalies associated with El Ninõ followed by a La Ninã. The monthly averaging effectively filters out fast Kelvin waves from the picture leaving only gradual interannual changes. It has been argued that the Kelvin waves may have contributed to the exceptional strength of El Ninõ in 1997–98.


EQUATORIAL WAVES

EOF mode 2

20

20

1

10

10

0.5

Latitude

Latitude

EOF mode 1

0

0

0

−10

−10 −20 150

200

285

−0.5 −1

−20

250

200

150

Longitude

250

Longitude Amplitude of the EOF modes

40 20 0 −20 −40 1982

1985

1987

1990 1992 Year

EOF mode 1

1995

1997

2000

EOF mode 2

Figure 16 The first two empirical orthogonal functions (EOFs) of the thermocline depth variations (approximated as the 20 1Cdegree isotherm depth) in the tropical Pacific. The upper panels denote spatial structure of the modes (nondimensionalized), while the lower panel shows mode amplitudes as a function of time (cf. Figure 12). The data are from hydrographic measurements combined with moored temperature measurements from the tropical atmosphere and ocean (TAO) array, prepared by Neville Smith’s group at the Australian Bureau of Meteorology Research Centre (BMRC). Adapted from Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Ninõ and La Ninã. Journal of Climate 13: 3551–3559.

the signal can interact remains to be seen, even though it has been argued that the Kelvin waves in Figure 15 may have contributed to the exceptional strength of El Ninõ of 1997–98. The observations also provide confirmation of the recharge oscillator mechanism, which can be demonstrated, for example, by calculating the empirical orthogonal functions (EOFs) of the thermocline depth (Figure 16). The first EOF (the left top panel) shows the spatial structure associated with changes in the slope of the thermocline, and its temporal variations are well correlated with SST fluctuations in the eastern equatorial Pacific, or the ENSO signal. The second EOF shows changes in the mean thermocline depth, that is, the ‘recharge’ of the equatorial thermocline. The time series for each EOF in the bottom panel of Figure 16 indicate that the second EOF (the thermocline recharge) leads the first EOF (a proxy for El Ninõ) by approximately 7 months.

Summary Early explanations of El Ninõ that relied on the straightforward generation of free Kelvin and Rossby waves by a WWB have been superseded. Modern theories consider ENSO in terms of a slow oceanic adjustment which occurs as a sum of continuously forced equatorial waves. The concept of ‘ocean memory’ based on the delayed ocean response to varying winds has become one of the cornerstones for explaining ENSO cyclicity. It is significant that although from the point of view of the ocean a superposition of forced equatorial waves is a direct response to the winds, from the point of view of the coupled ocean–atmosphere system it is a part of a natural mode of oscillation made possible by ocean– atmosphere interactions. Despite considerable observational and theoretical advances over the past few decades many issues are still being debated and each El Ninõ still brings


15 Nov. 98

12 Jul. 98

286

EQUATORIAL WAVES

5° N 0 5° S

5° N 0 5° S 140° E

180° E

140° W 22 °C

100° W 24 °C

26 °C

60° W

20° W

20° E

28 °C

Figure 17 Tropical instability waves (TIWs): 3-day composite-average maps from satellite microwave SST observations for the periods 11–13 July 1998 (upper) and 14–16 November 1998 (lower). Black areas represent land or rain contamination. The waves propagate westward at approximately 0.5 m s 1. From Chelton DB, Wentz J, Gentemann CL, de Szoeka RA, and Schlax MG (2000) Satellite microwave SST observations of trans-equatorial tropical instability waves. Geophysical Research Letters 27(9): 1239–1242.

surprises. The prolonged persistence of warm conditions in the early 1990s was as unexpected as the exceptional intensity of El Ninõ in 1982 and again in 1997. Prediction of El Ninõ also remains problematic. Not uncommonly, when a strong Kelvin wave crosses the Pacific and leaves a transient warming of 1–2 1C in the eastern part of the basin, a question arises whether this might be a beginning of the next El Ninõ. To what degree, random transient disturbances influence ENSO dynamics remains unclear. Many theoretical and numerical studies argue that high-frequency atmospheric disturbances, such as WWBs that excite Kelvin waves, can potentially interfere with ENSO and can cause significant fluctuations in its period, amplitude, and phase. Other studies, however, insist that external to the system atmospheric ‘noise’ has only a marginal impact on ENSO. To resolve this issue we need to know how unstable the coupled system is. If it is strongly damped, there is no connection between separate warm events, and strong wind bursts are needed to start El Ninõ. If the system is sufficiently unstable then a self-sustained oscillation is possible. The truth is probably somewhere in between – the coupled system may be close to neutral stability, perhaps weakly damped. Random atmospheric disturbances are necessary to sustain a quasi-periodic, albeit irregular oscillation. Another source of random perturbations that affects both the mean state and interannual climate variations is the tropical instability waves (TIWs) typically observed in the high-resolution snapshots of tropical SSTs (Figure 17). These waves, propagating westward with typical phase speed of roughly 0.5 m s 1, are excited by instabilities of the zonal equatorial currents with strong vertical and

horizontal shear. The wave dynamical structure corresponds to that of cyclonic and anticyclonic eddies having maximum velocities near the ocean surface and penetrating into the ocean by a few hundred meters. The waves can affect the temperature of the equatorial cold tongue, and the properties of ENSO, by modulating meridional heat transport from the equatorial Pacific. Overall, the role of the TIWs remains a subject of intensive research which includes the effect of these waves on the coupling between the wind stress and SSTs and the interaction between the TIWs and Rossby waves.

See also El Ninõ Southern Oscillation (ENSO). El Ninõ Southern Oscillation (ENSO) Models. Pacific Ocean Equatorial Currents.

Further Reading Battisti DS (1988) The dynamics and thermodynamics of a warming event in a coupled tropical atmosphere/ ocean model. Journal of Atmospheric Sciences 45: 2889--2919. Chang PT, Yamagata P, Schopf SK, et al. (2006) Climate fluctuations of tropical coupled system – the role of ocean dynamics. Journal of Climate 19(20): 5122–5174. Chelton DB and Schlax MG (1996) Global observations of oceanic Rossby waves. Science 272(5259): 234--238. Chelton DB, Schlax MG, Lyman JM, and Johnson GC (2003) Equatorially trapped Rossby waves in the presence of meridionally sheared baroclinic flow in the Pacific Ocean. Progress in Oceanography 56: 323--380.


EQUATORIAL WAVES

Chelton DB, Wentz J, Gentemann CL, de Szoeka RA, and Schlax MG (2000) Satellite microwave SST observations of trans-equatorial tropical instability waves. Geophysical Research Letters 27(9): 1239–1242. Fedorov AV and Philander SG (2000) Is El Ninõ changing? Science 288: 1997--2002. Fedorov AV and Philander SG (2001) A stability analysis of tropical ocean–atmosphere interactions: Bridging measurements and theory for El Ninõ. Journal of Climate 14(14): 3086–3101. Gill AE (1982) Atmosphere-Ocean Dynamics, 664p. New York: Academic Press. Jin FF (1997) An equatorial ocean recharge paradigm for ENSO. 1. Conceptual model. Journal of the Atmospheric Sciences 54: 811--829. Kessler WS (2005) Intraseasonal variability in the oceans. In: Lau WKM and Waliser DE (eds.) Intraseasonal variability of the Atmosphere-Ocean System, pp. 175--222. Chichester: Praxis Publishing. Levitus S and Boyer T (1994) World Ocean Atlas 1994, Vol. 4: Temperature NOAA Atlas NESDIS4. Washington, DC: US Government Printing Office. McPhaden MJ (1999) Genesis and evolution of the 1997– 98 El Ninõ. Science 283: 950–954.

287

Meinen CS and McPhaden MJ (2000) Observations of warm water volume changes in the equatorial Pacific and their relationship to El Ninõ and La Ninã. Journal of Climate 13: 3551--3559. Philander G (1990) El Ninõ, La Ninã, and the Southern Oscillation. International Geophysics Series, 293p. New York: Academic Press. Schopf PS and Suarez MJ (1988) Vacillations in a coupled ocean atmosphere model. Journal of the Atmospheric Sciences 45: 549--566. Wang C, Xie SP, and Carton JA (2004) Earth’s climate: The ocean–atmosphere interaction. Geophysical Monograph 147, American Geophysical Union, 405p. Zebiak SE and Cane MA (1987) A model El Ninõ-southern oscillation. Monthly Weather Review 115(10): 2262--2278.

Relevant Website http://www.pmel.noaa.gov – Tropical Atmosphere Ocean Project, NOAA.


ESTIMATES OF MIXING J. M. Klymak, University of Victoria, Victoria, BC, Canada J. D. Nash, Oregon State University, Corvallis, Oregon, OR, USA & 2009 Elsevier Ltd. All rights reserved.

Introduction Mixing in the ocean redistributes tracers, driving physical and biogeochemical dynamics. The mixing of the ‘active’ tracers, temperature and salinity, changes the density of seawater, creating pressure gradients that can drive mean currents. For example, in overturning circulations that range in scale from small estuaries to the global ocean, the mixing of buoyant fluid through the interior sets the conversion rate of potential energy and directly controls the strength of overturning. Momentum is also diffused by turbulent mixing, which transmits forces from the ocean surface and boundaries into the interior. The mixing of ‘passive’ scalars, such as nutrients, carbon dioxide, and oxygen, is important to understanding biological cycles in the ocean. Phytoplankton rely on vertical mixing to transport recycled nutrients into the sunlit near-surface waters. The mixing of carbon dioxide ultimately affects its storage in the ocean and removal from the atmosphere.

0.54

Most of the mixing in the interior of the ocean is thought to take place when internal waves break due to convective or shear instabilities. A numerical simulation serves to illustrate a typical ocean mixing event (Figure 1). A Kelvin–Helmholtz billow is triggered on an interface between warm and cold water when the warm water moves to the right faster than the cold water. A wave-like instability grows and two vortices form (Figure 1(a)). The vortices pair to create a breaking vortex (Figure 1(b)). Further instabilities ensue, creating a fully turbulent and three-dimensional flow field (Figure 1(c)) that decays as small-scale shear and temperature variability if mixed away. Breaking events like this are believed to dominate mixing in the ocean. They dominate because molecular diffusivity acting on ‘large-scale’ gradients (tens of meters) is very ineffective at mixing. Mixing is ultimately accomplished by molecular processes via Fickian diffusion, that is, the irreversible flux of property C is proportional to its three-dimensional gradient and the molecular diffusion coefficient kC: f C ¼ kC rC

½1

For temperature, a thermodynamic tracer, kTE10 7 m2 s 1 and for salinity and other tracers kSE10 9 m2 s 1. At large scales, representing the nonturbulent flow, gradients are small and the molecular flux is slow. In the absence of turbulence, a spike of

(a)

(b)

(c)

(d)

Z (m)

0.27 0.00 −0.27 −0.54 0.5 0.54

0.0

0.00

/o

Z (m)

0.27

−0.27 −0.54

−0.5

Figure 1 A numerical simulation of turbulent mixing (Smyth et al., 2001). The event is triggered by a shear instability between an upper layer of warm water (red) moving to the right and a lower layer of cold water (blue) moving to the left. The initial pair of vortices (a) combine to create a single large breaking event (b). This becomes fully turbulent and three dimensional (c) at which point there is large irreversible diffusion of heat. Diffusion continues until a large volume of mixed fluid results (d).

288


ESTIMATES OF MIXING

high temperature introduced into an otherwise isothermal fluid would take over 4 months to spread just 1 m through molecular diffusion alone. However, the stirring driven by the breaking of ‘fine-scale’ (order 1–10 m) waves maintains gradients at the ‘microscale’ (order 1 mm to 1 cm). The microscale gradients can be very large, Figures 1(b) and 1(c), driving a larger molecular flux that mixes large-scale gradients more efficiently. Typical ocean turbulence diffusivities are at least 100 times greater than molecular diffusivities, and can diffuse the above temperature spike over 1 m in O(1 day). It is useful to parametrize the turbulent stirring that drives the turbulent flux in terms of gradients of the mean fields, C. We do this by defining a ‘turbulent diffusivity’ KC so that FC E KC rC

½2

Note that this parametrization has the same form as eqn [1], so we are drawing a direct analogy between the random walk that accomplishes mixing on the microscale and the stirring that takes place on the finescale of a breaking wave. The power of this concept is that the random walk of the stirring is on ‘large’ scales, and does not change for different tracers in the water. It is generally assumed that all variance created via stirring at large scales is ultimately transformed to small enough scales where it diffuses via molecular processes. Thus, an estimate of the turbulent diffusivity for one tracer may be applied to other tracers experiencing the same turbulent flow. Therefore we drop the subscript and discuss the turbulent diffusivity, K, as a dynamic property of the flow. The methods discussed below find that the turbulent diffusivity in the open ocean is KE100–1000 kT, and much more in shallow water and near topography. The accumulated effect of this mixing becomes an important term in understanding the circulation of the oceans.

Approaches to Quantifying Mixing The Advection–Diffusion Balance

The sequence of events shown in Figure 1 illustrates the different methods of how we quantify mixing in the ocean. Suppose we are interested in the mixing of temperature T in a fluid. It is often assumed that one can separate the scales of turbulent motions from those of the mean flow, allowing one to write: T ¼ T þ T0

½3

u ¼ u þ u0

½4

289

where the primes represent the ‘turbulent’ part of the flow and the overbars the mean quantities. In Figure 1, the horizontally averaged velocity and temperature represent the mean, and deviations from these the turbulence. This so-called ‘Reynolds decomposition’ allows us to transform the advection– diffusion equation qT/qt þ u rT ¼ kTr2T into an evolution equation for the mean temperature T: DT ¼ kT r2 T þ r hu0 T 0 i Dt

½5

¼ r ðf T þ FT Þ

½6

Here, the material derivative D=Dt ¼ ð@=@t þ u rÞ is the change in time following a parcel in the mean flow, and the angle brackets denote an average in time and space over a turbulent event. Equation [5] shows that T depends not only on T and u, but also on the correlation /u0 T 0 S which we term the turbulent heat flux, often approximated using the Fickian analogy [2] as: FT ¼ hu0 T 0 iE KrT

½7

In most places in the ocean, the turbulent flux acting on the mean gradients is much greater than the molecular one, |FT|c|fT|, implying that we can drop the first term on the right-hand side of eqn [5]. Considering Figure 1 again, suppose we integrate eqn [5] over a volume, v, defined by the lower half of the domain, with z ¼ 0 the top of the volume. There are no fluxes out the sides or bottom of the volume, and there is no mean flux out the top ðwðz ¼ 0Þ ¼ 0Þ. The only flux of temperature is the turbulent one through z ¼ 0, so that the change of temperature in the volume can be calculated by @ @t

Z

T dV ¼ V

I

hw0 T 0 i dA

½8

A

where A is the surface at z ¼ 0. In Figure 1, there is an increase in the mean temperature of the volume, so the left-hand side of eqn [8] is greater than zero. The tendrils that drop below z ¼ 0 are warmer than the mean, so T 0 40, and they are moving down, so w0 o0, therefore w0 T 0 o0. If the tendrils are completely diffused away by molecular mixing in the volume, then warm water will have been left behind and the temperature in the volume will increase. The real situation is more complicated. Tendrils are further strained and stretched, and some of the warm water rises again out of the volume. However, on average some is always exchanged so that /w0 T 0 So0 and net warming takes place in the lower volume. Note that there is an equal amount of cooling in the upper volume.


290

ESTIMATES OF MIXING

The Gradient–Variance Balance

In order for stirring to be irreversible, the gradients produced must be diffused away by molecular processes. For temperature, this is described formally through the evolution equation for turbulent temperature gradient variance |rT 0 |2. Temperaturegradient variance is a nonintuitive quantity to consider, but it is the best measure of ‘stirring’, and is related to the thermodynamic quantity of entropy. For steady state, homogeneous turbulence it can be shown that: D E 2 ½9 hu0 T 0 i rTEkT jrT 0 j This states that the net production of gradient variance by turbulent velocities is balanced by its destruction by molecular diffusion (there are transport terms that have been dropped, hence the approximation). The averages represented by the angle brackets must be collected over long enough time that the irreversible part of the ‘turbulent’ flux is measured. The rate of destruction of the turbulent gradients is fundamental to quantifying mixing in the ocean and is written as D E 2 ½10 w 2kT jrT 0 j

Large-Scale Estimates Large-scale estimates are made on quantities measured on vertical scales greater than a meter. They are based on estimating the mixing terms in the advection–diffusion equation of the tracer C: @C þ u rC ¼ r ðKrCÞ @t

½11

where again, K is the turbulent diffusivity. Often the right-hand side is replaced by @=@zðK@C=@zÞ because mean vertical gradients exceed horizontal ones, further reducing to K@ 2 C=@z2 for spatially uniform K.

Tracer Budgets (Inverse Methods)

The rate of mixing can be estimated from an integrated version of the mean tracer equation (eqn [11]) if we constrain a volume of water and assume its contents are in steady state. A concrete example of the budget method is from data collected in the Brazil Basin in the Southwest Atlantic. Here, water colder than 1 1C produced in the Antarctic flows north into the basin through the Vema Channel (Figure 2). No water that cold is observed to leave the basin, therefore the incoming water must be warmed by mixing before it leaves. Quantitatively, the mixing estimate is made by integrating the remaining terms in eqn [11] over the volume:

The first and conceptually simplest method to measure mixing is to release a man-made tracer and measure its vertical spread over time (see Tracer Release Experiments) . Briefly, if we follow the parcel of water and assume a constant vertical diffusivity K, then the spread of the tracer, C, is governed by the diffusion equation ½12

Adv: top

in Vema

Turb: top

zfflfflfflfflfflfflffl}|fflfflfflfflfflfflffl{ zfflfflfflfflfflfflffl I ffl}|fflfflfflfflfflfflfflffl{ zfflfflfflfflffl I ffl}|fflfflfflfflfflffl{ I uT dA1 wT s dAS ¼ FT dAs 1

Purposeful Tracer Releases

@C @2C ¼K 2 @t @z

If the tracer is injected as a spatial delta function at t ¼ 0, then the solution is an ever-widening vertical cloud described by a Gaussian. The larger the K, the faster the cloud spreads. Direct tracer releases are elegant and definitive in their results. There are no confounding sources or sinks of the dye in the ocean, so tracking the vertical spread is unambiguous. (Please note that care should be taken interpreting the terms ‘vertical’ and ‘horizontal’ or ‘lateral’. By these terms we really mean perpendicular and parallel to surfaces of constant density respectively, or ‘diapycnal’ and ‘isopycnal’. The distinction is conceptually important, but notationally less convenient, so we use the shorthand here.) However the technique is difficult to perform, limiting its routine use. Specialized equipment and analysis methods are needed to release the dye and then analyze the water samples to find minute quantities of tracer in the water. Furthermore, horizontal stirring and advection of a dye patch can spread it horizontally to such an extent that it is very difficult to find all the dye using finite ship resources.

S

½13

S

We also know that there is just as much water entering the volume as leaving through the upper surface, so QS ¼ Q1, where QS ¼

I

w dAS

½14

S

is the volume flux through the upper surface, and Q1 is similarly defined and is the flux through Vema Channel. Only the advective transport of heat into


ESTIMATES OF MIXING

s

FT dAs

291

Ts uT dA1

Ts

w dAs s

Figure 2 Sketch of heat fluxes in and out of a volume bounded in the vertical by an isotherm Ts, and at a strait by the dashed line.

the basin needs to be measured (the left-hand side of [13]) to determine the turbulent heat flux through the upper bounding surface. Using a combination of moorings and shipboard cruises, Hogg et al. estimated these transports by assuming that the deviations from the mean temperature and velocity values entering Vema Channel are uncorrelated: I

uT dA1 EQ T 1

½15

1

They then defined the upper surface as the TS ¼ 1 1C isotherm and assumed a constant vertical temperature gradient along that surface, allowing the mean turbulent diffusivity at that surface to be determined as

difficult amount of change to detect in open ocean basins with any confidence. Furthermore, the velocities and tracers measured at the boundaries of the volume must be well constrained and shown to be in steady state. This is very difficult as velocities and tracers are estimated from a few individual ship tracks that each take a month or more to complete, often months or years apart. Not surprisingly, the most convincing inverse estimates have come from well-constrained topographies like the Brazil Basin where the flow and temperatures into the basin can be monitored with a few long-term moorings in the channel, and the bounding isotherm can be mapped with a high degree of confidence.

Fine- and Microscale Estimates

K¼

@T Q Ts T1 As @z

1 ½16

6 3 1 For the Brazil Basin, Q ¼ 3.7 10 m s , AS ¼ 5 12 2 10 m , T 1 ¼ 0:35 1C, and the mean temperature gradient at the 11 C surface is @T/@zE 2 10 3 1C m 1. The mean turbulent diffusivity across this interface is therefore KE2:5 104 m2 s1 . Inverse estimates of mixing are routinely made from hydrographic sections in the ocean, where the same concepts are used to find the flux of heat and density at different depths in a series of volumes. Often many vertical layers and sections are used. If well constrained, this method of estimating the heat flux (and thus the diapycnal mixing) would be unambiguous in providing basin-average estimates. A one-dimensional advection–diffusion balance can be used to estimate the turbulent diffusivity in the deep ocean from large-scale tracer profiles. This simple inverse estimate finds average turbulent diffusivities of K ¼ 10 4 m2 s 1 in the North Pacific. The principal difficulty with the inverse method is the assumption that the system is in steady state. Inverse estimates in the open ocean indicate vertical velocities of a couple of meters per year. This is a

Fine- and microscale measurements estimate turbulent stirring or mixing by directly observing the turbulence. These methods have the advantage over budget-based estimates in that they also elucidate what causes the turbulence. However, these methods require specialized instrumentation as the sensors used to measure microscale quantities must be small, respond quickly, and be capable of recording a very large dynamic range. In addition, the vehicles they are mounted on must suppress vibrations as much as possible to prevent contamination of the small-scale signals (Figure 3). Vibration and spatial resolution concerns limit the speed with which these profilers can be dropped or towed as well. In the following, we describe a series of methods that (1) directly measure the turbulent stirring of a fluid using the eddy correlation technique, (2) directly measure the molecular destruction of temperature gradients, (3) estimate the mixing by relating the buoyancy flux to the energetics of the turbulence. Finally, we describe two techniques that enable mixing to be estimated from large-scale measurements of (1) statically unstable fluid (Thorpe scales) and (2) energy in the internal wave field (the Gregg– Henyey method).


292

ESTIMATES OF MIXING

Figure 3 Microstructure platforms that the authors have worked with. In all cases the sensors are on the nose of the vehicles. The upper three are profilers. Left to right: Absolute velocity profiler (University of Washington), advanced microstructure profiler (University of Washington) and Chameleon (Oregon State University). The lower instrument, Marlin (Oregon State University), is towed.

Direct Eddy Correlation

With careful and specialized measurements it is possible to estimate mixing by quantifying the stirring of the fluid. As described above, this means quantifying the stirring of cold water upward into warm water by measuring vertical velocity fluctuations, w0, and temperature fluctuations, T 0. One of the few attempts to apply this method in the open ocean demonstrates its difficulty (Figure 4). Temperature and velocity were acquired along a horizontal path using a towed instrument outfitted with thermistors and shear probes. While the raw signals are large and very active (Figure 4(a)), the product w0 T 0 is not one-sided. Instead, it has instantaneous values that are large and can be of either positive and negative sign, such that the fluctuations are far greater than the mean correlation /w0 T 0 S. This is a general problem since turbulence is sporadic, and stirring is both downgradient (i.e., transports heat from regions of warm fluid) and upgradient (i.e., transports heat to regions of warm fluid), the latter representing restratification of partially mixed fluid. Since much of w0 T 0 is reversible (i.e., just stirring fluid

that is not immediately mixed), the eddy correlation technique must be made over long times to produce stable estimates of the irreversible flux. In this case, the background vertical gradient dT=dz is positive, so /w0 T 0 So0 represents a downgradient flux, as appears most frequently in Figure 4(b). This method does not enjoy wide use. Determining what is ‘mean’ and what is ‘turbulent’ is very difficult from the limited measurements possible with a vertical or horizontal profiler. The data presented here were simply bandpassed, with large scale motions considered to be nonturbulent. However choosing what is ‘large scale’ requires some art. A second difficulty is estimating the vertical velocities in the ocean. In this instance, the vertical velocities were w0 E0.03 m s 1, large enough that the method was deemed possible. The final limitation is gathering enough statistics of the turbulence to make estimates of mean fluxes. Microscalars (Osborn–Cox Method)

In contrast to measuring the fine-scale stirring of the tracer (i.e., /w0 T 0 S above), the Osborn–Cox method


ESTIMATES OF MIXING

(a)

293

Tope S7 Blocks: 2197-2244 0.08

0.03 0.02

0 −0.04

0.01

−0.08

w (m s−1)

(°C)

0.04

0

−0.12

−0.01

−0.16 0

5

10

15

20

25

30

35

40

45

50

55

60

65

70

75

0

5

10

15

20

25

30

35 40 x (m)

45

50

55

60

65

70

75

w (°C m s−1) × 104

(b)

1 0

1

Figure 4 (a) Spatial series of high-passed temperature and vertical velocity from a towed vehicle in the upper ocean thermocline (Fleury and Lueck, 1994). The high-passing procedure is meant to emphasize turbulent fluctuations. (b) The correlation between these observations.

quantifies the rate of molecular diffusion of scalar variance at the microscale. By considering the evolution equation for microscale scalar variance, the turbulent diffusivity K is related to the rate of destruction of scalar variance w (eqn [10]). Because this method measures the rate of irreversible molecular mixing, it is one of the most direct measures of quantifying K. The most commonly measured scalar is temperature, as it is relatively easy to measure and as it diffuses at larger scales (i.e., 1 mm to 1 cm) than chemical constituents like salt. To measure such scales, a sensor must be small and respond very rapidly. Microbead thermistors – coated with a thin film of glass to electrically insulate them from seawater (Figure 5) – are generally used for this purpose. This sensor yields high-resolution temperature gradients such as those shown in Figure 6(b). Only one dimension of the gradient is measured, so we estimate w as *

+ @T 0 2 ½17 w ¼ 6kT @z where we have assumed that the turbulence is isotropic (i.e., the variance of the gradients is the same

in all directions). The turbulent diffusivity simply relates the intensity of small-scale gradients to the large-scale temperature gradient as D E ð@T 0 =@zÞ2 w ½18 K ¼ 3kT 2 ¼ 2 @T=@z 2 @T=@z Like other methods, there are a number of limitations. Temperature probes require a finite amount of time to diffuse heat through their insulation and the thermal boundary that develops in the surrounding seawater. This smooths the signals and coarsens the measurement. If probes can be lowered slowly enough to allow heat to diffuse through the coating, but fast enough to capture a synoptic snapshot of the turbulent event, then all of the gradient variance could be resolved and w measured. However, the required 10–20 cm s 1 profiling speed reduces the number of realizations that may be captured, so there is a trade-off between resolution and statistics. In practice, most sensors are deployed too rapidly and not all of the variance is measured. Corrections can be applied by fitting data to a universal spectrum (i.e., the ‘Batchelor’ or ‘Kraichnan’ spectrum) which allows extrapolation of resolved


294

ESTIMATES OF MIXING

0.25 mm

2

6.4-mm Stainless Hard diameter steel tube epoxy

0.15mm

3

1

4

Rubber tip

5

Bimorph Heat shrink beam tubing

Electrical leads

Figure 5 Three common sensors used on microstructure instruments. Left: A glass-bead thermistor (Gregg, 1999); upper scale in mm. Center: a four-electrode microconductivity probe (Nash and Moum, 1999). 1–4 are electrodes and 5 is an insulating glass. Right: Schematic of a piezoelectric shear probe (Gregg, 1999).

(a) T (°C)

5.12 5.11

(b)

0.5

T′

5.1

0

* v /* x (s1)

(c)

0.5 0.1 0 0.1

* w /* x (s1)

(d)

0.1 0 0.1 100

105

110

115

120

125 Time (s)

130

135

140

145

150

Figure 6 Microscale data collected using the towed vehicle Marlin in the open ocean (Moum et al., 2002). This data was collected with the instrument moving at 1 m s 1, so time is equivalent to distance in m. (a) Temperature, (b) temperature gradient, (c) and (d) velocity gradients (shear). Note that where there is high velocity variance there is usually high temperature variance.

measurements to higher wavenumbers (Figure 7). Unfortunately, the spectral levels and wavenumber extent of the spectrum both depend on the turbulent kinetic energy dissipation rate e, so an independent measure of microscale shear variance (see below) is required to accurately apply such corrections. It is also possible to use microconductivity probes to measure temperature on spatial scales of 10 3 m (Figure 5, center panel). Conductivity is a rapid measurement, so speed through the water does not limit probe response, only the physical configuration of the probes. The difficulty with this measurement is

that conductivity depends on both the temperature and salinity of seawater. Thus, microconductivity works best for determining the mixing rate of temperature in water that has small salinity variations. The last difficulty, shared with the other microscale methods, is that turbulence is intermittent, so a large number of samples need to be made in order to characterize the turbulence level of a given locale. However, unlike the estimate /w0 T 0 S, w is a direct measure of irreversible mixing, so does not need multiple realizations of the same turbulent event to converge.


ESTIMATES OF MIXING

dT/dz ((K m−1)cpm−1)

10−3

295

10−3 Weak turbulence

Strong turbulence Response 10−4

10−4

10−5

10−5 raw

fit 10−6 10−1

100

10−6 10−1

101

100

101

kx (cpm) Figure 7 Fit of universal turbulence spectra (red curves) to data spectra collected near Hawaii (Klymak and Moum, 2007). The raw signals (gray) have been corrected (black) for the temporal response of the thermistor. Note that noise at high wavenumbers may contaminate the spectra of weak turbulence.

Estimates from Energy Considerations (Osborn Method)

The most common method of estimating ocean mixing rates is based on quantifying the energetics of the turbulence, not the mixing itself. It is widely employed because the energetics can be measured from rapidly profiling sensors, and because energy measurements are useful in their own right. This method was originally proposed by Osborn in 1980, and is based on the observation that temperature-gradient variance occurs in accord with velocity gradient (shear) variance (i.e., compare Figures 6(b) and 6(c)). The argument is based on the local balance of turbulent kinetic energy. A turbulent event loses energy by viscous dissipation and by changing the background potential energy of the flow due to irreversible mixing. In a steady state, or time-averaged sense, this is expressed as P ¼ e þ Jb

½19

where P is the rate of production of turbulence by the mean flow due to some wave-breaking process, e is the rate of turbulent energy dissipation by viscosity, and Jb is the irreversible buoyancy flux due to mixing. The buoyancy flux is directly related to the turbulent mass flux Jb ¼ ghr0 w0 i=r where g is the gravitational acceleration, and r ¼ r þ r0 is the density. The method assumes that the turbulent buoyancy flux is a fixed ratio G of the dissipation:

unstratified water, the buoyancy flux must be zero by definition, but e can be substantial. However, observations indicate that for much of the stratified ocean GE0.2 7 0.05. (Note that G can be related to the ‘mixing efficiency’ Rf ¼ Jb/P ¼ G/(G þ 1) via eqn [19]). Measurements of the dissipation rate of turbulent kinetic energy e are made from microstructure profilers (Figure 3) in much the same way that measurements of w are made. A small shear probe (Figure 5, right panel) measures velocity shear (Figures 6(c) and 6(d)). The shear spectrum is calculated and integrated to get an estimate of the dissipation rate of turbulent kinetic energy e, again using universal spectra as a guide. Fortunately, both the spectral amplitude and wavenumber extent of shear spectra scale with e, so that universal spectra may be fit to a limited range of wavenumber range of the shear spectrum, avoiding poorly resolved wavenumbers. But unlike the measurement of w using thermistors, measurement of turbulent shear variance is easily contaminated by the slightest vibration of the measurement platform. This necessitates the use of specialized profilers that minimize coherent eddy shedding and decouple ship motion from the sensor (Figure 3). Just like temperature, the flux of density can be parametrized by a turbulent diffusivity so that Jb ¼

g @r K ¼ KN 2 r0 @z

for stratification Jb EGe

½20

This is a somewhat bold assumption as obvious counterexamples can be found. For instance, in


g N ¼ r0 2

!

dr dz

!

½21

296

ESTIMATES OF MIXING

Combining, we use measurements of e to estimate e K¼G 2 N

½22

This method has greatly increased the number of estimates of mixing in the ocean. Microstructure profilers have been deployed in a wide array of environments: in the open ocean, over rough and abrupt topography, and in coastal waters, giving us a large variety of environments in which ocean mixing has been estimated. As tenuous as it is, the assumptions used in this method are mitigated by the fact that we know that dissipation rates vary by orders of magnitude throughout the ocean so that the distribution of dissipation roughly mimics the distribution of mixing, except in well-mixed regions. Thorpe scales The dissipation rate can also be estimated from less-specialized instruments. Breaking internal waves, like that shown in Figure 1, lift

p (MPa)

(a)

dense water above light water. The dissipation rate of the water that goes into turbulence from this uplift can be estimated from the size of the overturn LT : eE0:64L2T N 3

½23

This has been shown to give unbiased estimates of the dissipation rate if enough profiles are collected. Note that, at open ocean dissipation rates and stratifications, LT is quite small and the small density differences make detecting overturns subject to noise constraints. A coastal example is shown in Figure 8, where braids indicative of shear-instabilities drive density overturns with LT E 5 m. The overturns coincide with strong turbulence. There is also a strong correspondence of e and w in these data. This study compared the two estimates of K from these separate microstructure estimates and found similar results.

100 300 500 700 900 1100 1300 1500 Distance (m) + + + + + + + + + ++ + + + + + + + + ++ + + + + + + + + + + + AMP # 51 52 53 54 55 56 57 58 −72 0.2 23.5 VSS (dB) 0.4 23.75 −77 24.0 0.6 24.25

−82

0.8 (b) 0.2 p (MPa)

23.5 0.4

log (K2 s−1)

0.6

23.75

24.0 24.25

p (MPa)

(c) 0.8 0.2 0.4

−10

23.5 0

10

0.6

24.0

23.75

Lt (m)

24.25

(d) 0.8 0.2 p (MPa)

23.5 0.4

0.8

23.75

24.0

0.6

log (W kg−1)

24.25 0

0.2

0.4 Time (h)

0.6

0.8

Figure 8 An example of a deterministic turbulent event measured four ways. This is a shear instability, similar in dynamics to the instability in Figure 1, observed in Admiralty Inlet, Washington (Seim and Gregg, 1994). (a) Acoustic backscatter from turbulence microstructure. This visualizes the braids between 0.2 and 0.6 h. Vertical white lines are microstructure profiler drops, horizontal white lines are contours of r. (b) w estimated from profiler. (c) Density overturns measured with the profiler. Note how there are large overturns associated with the braids. (d) The turbulence dissipation rate, e estimated from shear probes. For all the panels, 0.1 MPa ¼ 10 dbar ¼ 10-m depth.


ESTIMATES OF MIXING

−3

297

log(K ) m2 s−2

1000 −4

6000

0

2000

Hawaiian Rise

5000

−5

Sitito Iozima Ridge

4000

Japan

3000

Caroline Ridge

Admiralty Islands Ridge

z (m)

2000

4000

6000

GM IW

−6

8000

10 000

12 000

14 000 Moonless Smts

Isu Trench

Nankei Trench

Papua New Guinea

r (km)

Figure 9 Indirect estimate of mixing from the western Pacific Ocean using large-scale oceanic data (Kunze et al., 2006). Internal wave energy levels were used to estimate e, and hence K.

Gregg–Henyey method The Osborn method may be extended further using the observation that away from boundaries and strongly sheared currents the dissipation rate is directly related to the energy in the internal wave field. Models have been developed that estimate the rate at which energy cascades through a steady-state internal wave field. It is easier to estimate the energy of the wave field with finescale sensors than it is to directly measure the microscale. For instance, these methods have allowed the estimate of mixing using routine hydrographic data of the world oceans (Figure 9). The internal-wave energy method has limited application in regions where the internal wave field is not in equilibrium with external forcing, in particular near topography, or where turbulence is generated by noninternal wave process at boundaries. Somewhat frustratingly, this is where the dissipation rates are the strongest.

microscale methods do not agree as well. Inverse methods indicate turbulent diffusivities on the order of K ¼ 10 4 m2 s 1, while microstructure measurements are challenged to find average turbulent diffusivities this high. Recent attention has been directed toward finding enhanced mixing near boundaries. This zeroth-order problem will continue to require much effort and ingenuity to solve. Higher-order testing of the assumptions that go into these measurements are ongoing, aided by innovations in measurements and numerical methods.

Summary

e

Substantial effort has gone into estimating the rate of mixing in the ocean. The problem is hard to tackle directly, so great ingenuity has been used to devise indirect methods of making the observations. Mixing measurements have been made in many environments in the open and coastal ocean, lakes, and estuaries. In the Brazil Basin, where the source of deep water is well constrained, estimates of K from the basin-scale estimates agree quite well with microstructure estimates. In the open ocean, however, large-scale and

Nomenclature fC FC Jb K N G

kC w

molecular flux of scalar C turbulent flux of scalar C turbulent buoyancy flux turbulent diffusivity buoyancy frequency ratio of buoyancy flux to viscous dissipation rate of turbulent kinetic energy dissipation molecular diffusivity for scalar C rate of temperature variance dissipation

See also Energetics of Ocean Mixing. Internal Tides. Inverse Modeling of Tracers and Nutrients. Long-Term Tracer Changes. Storm Surges. Three-Dimensional (3D) Turbulence. Tracer Release Experiments.


298

ESTIMATES OF MIXING

Tracers of Ocean Productivity. Upper Ocean Mixing Processes. Wind- and Buoyancy-Forced Upper Ocean.

Further Reading Dillon TM (1982) Vertical overturns: A comparison of Thorpe and Ozmidov length scales. Journal of Geophysical Research 87: 9601--9613. Eckart C (1948) An analysis of the stirring and mixing processes in incompressible fluids. Journal of Marine Research 7: 265--275. Fleury M and Lueck R (1994) Direct heat flux estimates using a towed vehicle. Journal of Physical Oceanography 24: 810--818. Ganachaud A and Wunsch C (2000) Improved estimates of global ocean circulation, heat transport and mixing from hydrographic data. Nature 408: 453--457. Gregg MC (1989) Scaling turbulent dissipation in the thermocline. Journal of Geophysical Research 94: 9686--9698. Gregg MC (1999) Uncertainties in measuring e and wt. Journal of Atmospheric and Oceanic Technology 16: 1483--1490. Henyey FS, Wright J, and Flatte´ SM (1986) Energy and action flow through the internal wave field. Journal of Geophysical Research 91: 8487--8495. Hogg N, Biscaye P, Gardner W, and Schmitz WJ, Jr. (1982) On the transport and modification of Antarctic bottom water in the Vema Channel. Journal of Marine Research 40: 231--263. Johnson HL and Garrett C (2004) Effects of noise on Thorpe scales and run lengths. Journal of Physical Oceanography 34: 2359--2373. Klymak JM and Moum JN (2007) Interpreting spectra of horizontal temperature gradients in the ocean. Part II: Turbulence. Journal of Physical Oceanography 37: 1232--1245. Klymak JM, Moum JN, Nash JD, et al. (2006) An estimate of tidal energy lost to turbulence at the Hawaiian Ridge. Journal of Physical Oceanography 36: 1148--1164. Klymak JM, Pinkel R, and Rainville L (2008). Direct breaking of the internal tide near topography: Kaena

Ridge, Hawaii. Journal of Physical Oceanography 38: 380–399. Kunze E, Firing E, Hummon JM, Chereskin TK, and Thurnherr AM (2006) Global abyssal mixing inferred from lowered ADCP shear and CTD strain profiles. Journal of Physical Oceanography 36: 1553--1576. Lumpkin R and Speer K (2003) Large-scale vertical and horizontal circulation in the North Atlantic Ocean. Journal of Physical Oceanography 33: 1902--1920. Moum JN (1996) Energy-containing scales of turbulence in the ocean thermocline. Journal of Geophysical Research 101: 14095--14109. Moum JN, Caldwell DR, Nash JD, and Gunderson GD (2002) Observations of boundary mixing over the continental slope. Journal of Physical Oceanography 32: 2113--2130. Nash J, Alford M, Kunze E, Martini K, and Kelley S (2007) Hotspots of deep ocean mixing on the Oregon continental slope. Geophysical Research Letters 34: L01605 (doi:10.1029/2006GL028170). Nash JD and Moum JN (1999) Estimating salinity variance dissipation rate from conductivity measurements. Journal of Atmospheric and Oceanic Technology 16: 263--274. Osborn TR (1980) Estimates of the local rate of vertical diffusion from dissipation measurements. Journal of Physical Oceanography 10: 83--89. Seim HE and Gregg MC (1994) Detailed observations of a naturally occurring shear instability. Journal of Geophysical Research 99: 10049--10073. Smyth WD, Moum JN, and Caldwell DR (2001) The efficiency of mixing in turbulent patches: Inferences from direct simulations and microstructure observations. Journal of Physical Oceanography 31: 1969--1992. St. Laurent LC, Toole JM, and Schmitt RW (2001) Buoyancy forcing by turbulence above rough topography in the abyssal Brazil Basin. Journal of Physical Oceanography 31: 3476--3495. Wunsch C and Ferrari R (2004) Vertical mixing, energy, and the general circulation of the oceans. Annual Review of Fluid Mechanics 36: 281--314.


ESTUARINE CIRCULATION K. Dyer, University of Plymouth, Plymouth, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 846–852, & 2001, Elsevier Ltd.

Introduction Estuaries are formed at the mouths of rivers where the fresh river water interacts and mixes with the salt water of the sea. Even though there is only about a 2% difference in density between the two water masses, the horizontal and vertical gradients in density causes the water circulation, and the mixing created by the tides, to be very variable in space and time, resulting in long residence times for pollutants and trapping of sedimentary particles. Both salinity and temperature affect density, but the salinity changes are normally of greatest influence. Most estuaries are the result of a dramatic rise in sea level of about 100 m during the last 10 000 years, following the end of the Pleistocene glaciation. The river valleys were flooded by the sea, and the valleys infilled with sediment to varying extents. The degree of infilling provides a wide range of topographic forms for the estuaries. Those estuaries that have had large inputs of sediment have been filled, and may have been built out into deltas, where the sediment flux is extreme. These are typical of tropical and monsoon areas. Where the sediment discharge was less, the estuary may still have many of the morphological attributes of river valleys: a sinuous, meandering outline, a triangular cross-sectional form with a deep central channel, and wide shallow flood plains. These are termed drowned river valleys, and are typical of higher latitudes. Where the land mass was previously covered by the glaciers, the river valleys may have been drastically overdeepened, and the U-shaped valleys became fiords. Because of the shallow water, the sheltered anchorages, and the ready access to the hinterland, estuaries have become the centers of habitation and of industrial development, and this has produced problems of pollution and environmental degradation. The solution to these problems requires an understanding of how the water flows and mixes, and how the sediment accumulates.

Definition Estuaries can be defined and classified in many ways, depending on whether one is a geologist/geographer,

a physicist, an engineer, or a biologist. The most comprehensive physical definition is that An estuary is a semi-enclosed coastal body of water that has free connection to the open sea, extending into the river as far as the limit of tidal influence, and within which sea water is measurably diluted with fresh water derived from land drainage.

There are normally three zones in an estuary: an outer zone where the salinity is close to that of the open sea, and the horizontal gradients are low; a middle zone where there is rapid change in the horizontal gradients and where mixing occurs; and an upper or riverine zone, where the water may be fresh throughout the tide despite there being a tidal rise and fall in water level.

Tides in Estuaries The tidal rise and fall of sea level is generated in the ocean and travels into the estuary, becoming modified by the shoreline and by shallow water. The tidal context is microtidal when the range is less than 2 m; mesotidal when the range is between 2 m and 4 m; macrotidal when it is between 4 m and 6 m; and hypertidal when it is greater than 6 m. A narrowing of the estuary will cause the range of the tide to increase landward. However, this is counteracted by the friction of the seabed on the water flow, and the fact that some of the tidal energy is reflected back toward the sea. The result is that the time of high and low water is later at the head of the estuary than at the mouth, and the currents turn some tens of minutes after the maximum and minimum water level. This phase difference means that the tidal wave is a standing wave with a progressive component. In high tidal range estuaries, the currents and the range of the tide generally increase toward the head, until in the riverine section the river flow becomes important. These estuaries are often funnel-shaped with rapidly converging sides. The volume of water between high and low tide, known as the tidal prism, is large compared with the volume of water in the estuary at low tide. Since the speed of progression of the tidal wave increases with water depth, the flooding tide rises more quickly than the ebbing tide, leading to an asymmetrical tidal curve and currents on the flood tide larger than those on the ebb tide. These are flood-dominant estuaries. The high tidal range leads to high current velocities, which can


299

300

ESTUARINE CIRCULATION

produce considerable sediment transport, also dominant in the flood direction. When the estuary is relatively shallow compared with the tidal range, the wet cross-sectional area is greater at high water than at low water, and the discharge of water per unit of velocity is greater also. There is thus a greater discharge landward near high water than seaward around low water. This creates a Stokes Drift towards the head of the estuary that has to be compensated for by an extra increment of tidally averaged flow toward the mouth in the deeper areas. The Eulerian mean flow measured at a fixed point, the nontidal drift, is therefore greater than the flow due to the river discharge. Additionally, in the upper estuary where intertidal areas are extensive, the deeper channel changes from being flood-dominant to ebb dominance. This produces an obvious location for siltation and the need for dredging. In microtidal estuaries, friction exceeds the effects of convergence and the tidal range diminishes toward the head. These estuaries are generally fanshaped, with a narrow mouth and extensive shallow water areas. The currents at the mouth are larger than those inside, and the ebb current velocities are larger than the flood currents, i.e., the estuary would be ebb-dominant. It has been observed that the ratio of certain estuarine dimensions frequently appears to be constant, leading to the concept of an equilibrium estuary. In particular, the variations of breadth and depth along the estuary are often exponential in form. The O’Brien relationship shows that the tidal prism volume is related to the cross-sectional area of the estuary. As the tidal prism increases, the increased currents through the cross-section cause erosion and the area increases until equilibrium is reached. This implies that there is ultimately a balance between the amount of sediment carried landward on the flood tide and that carried seaward on the ebb, leading to zero net accumulation. This is a widely used concept for the prediction of morphological change.

classifications do not tell us much about the water and how its movement affects the discharge of pollutants and of sediment. The classifications, originally proposed by Pritchard in 1952, describe the tidally averaged differences in vertical and longitudinal salinity structure, and the mean water velocity profiles. These have been extended to consider the processes during the tide that contribute to the averages. Salt Wedge Estuaries

Where the river discharges into a microtidal estuary, the fresher river water tends to flow out on top of the slightly denser sea water that rests almost stationary on the seabed and forms a wedge of salt water penetrating toward the head of the estuary. The salt wedge has a very sharp salinity interface at the upper surface – a halocline. There is a certain amount of friction between the two layers and the velocity shear between the rapidly flowing surface layer and the almost stationary salt wedge produces small waves on the halocline; these can break when the velocity shear between the layers is large enough. The breaking waves inject some of the salty lower layer water into the upper layer, thus enhancing its salinity. This process, known as entrainment, is equivalent to an upward flow of salt water. It is not a very efficient process, but requires a small compensating inflow in the salt wedge. Salt wedge estuaries have almost fresh water on the surface throughout, and almost pure salt water near the bed. The slight amount of tidal movement of the water does not create much mixing, except in some of the shallower water areas, but does act to renew the salty water trapped in hollows on the bed. This process is shown in Figure 1. During the tide, the stratification remains high, but the halocline becomes eroded from its top surface during the ebb tide and rises during the flood tide because of new salt water intruding along the bottom. Fiords often have a shallow rock sill at their mouth that isolates the deeper interior basins. The renewal of bottom water is infrequent, often only every few years. Because fiords are so deep the lower layer is effectively stationary, and the mixing is dominated by entrainment.

Circulation Types From a morphological point of view, estuaries can be classified into many categories in terms of their shape and their geological development. These reflect the degree of infilling by sediment since the ice age. Present-day processes are also important in distinguishing those estuaries dominated by tidal currents and those whose mouths are drastically affected by the spits and bars created by wave-induced longshore drift of sediment. However,though important, these

Partially Mixed Estuaries

In mesotidal and macrotidal estuaries the tidal movement drives the whole water mass up and down the estuary each tide. The current velocities near to the seabed are large, producing turbulence, and creating mixing of the water column. This is a much more efficient process for mixing than entrainment, but it is greatest near the bed, whereas entrainment is maximum in mid-water. However, both processes


301


5

10

15

20

(A)

21

18

1 5 10 12 5

25

(A)

5 15

(B)

10

12

15

18

21

(B) Mean salinity

Mean velocity 0

S0

Mean velocity 0 us

Depth

Depth

Mean salinity

5 10 15 20 25

uf (C)

(C)

Figure 1 Diagrammatic longitudinal salinity and velocity distributions in a salt wedge estuary in PSU. (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strength of currents. (C) Tidally averaged mean salinity and velocity profiles.

can be active together. Turbulent mixing is a twoway process, mixing fresher water downward and salty water upward. As a result, the mean vertical profile of salinity has a more gentle halocline (Figure 2). The salt intrusion is now a much more dynamic feature, changing its structure regularly with a tidal periodicity. During the flood tide, the surface water travels faster up the estuary than the water nearer the bed, and the salinity difference between surface and bed is minimized. Conversely, during the ebb, the fresher surface water is carried over the near-bottom salty water, and the stratification is enhanced. This process is known as tidal straining. It is to be expected that entrainment will be large during the ebb tide, and turbulent mixing will be dominant on the flood tide. Since each tide must discharge an amount of fresh water equivalent to an increment of the river inflow, and the water involved in this is now salty, there must be a significant mean advection of salt water up the estuary near to the bed to compensate for that discharged. There is thus a mean outflow of water and salt on the surface, and a mean inflow near the bed. The latter must diminish toward the head of the salt intrusion, and there is a level of no net motion near to the halocline where the sense of the mean flow velocity changes. At the

ΔS

Figure 2 Diagrammatic longitudinal salinity and velocity distributions in a partially mixed estuary in PSU. (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strength of currents. (C) Tidally averaged mean salinity and velocity profiles, including the definition of parameters for Figure 4: S, salinity; DS, salinity difference normalized by S0 and circulation parameter us/uf; us, surface mean flow; uf, depth mean flow.

landward tip of the salt intrusion there is a convergence, a null zone, where the tidally averaged near-bed flow diminishes to zero. The vertical turbulent mixing acting in combination with the mean horizontal advection is known as vertical gravitational circulation. Because of the meandering shape of the estuary, the flow is not straight but tends to spiral on the bends. This leads to the development of lateral differences in salinity and velocity, the shallower regions generally being better mixed and ebb-dominated. Well-mixed Estuaries

When the tidal range is macrotidal or hypertidal, the turbulent mixing produced by the water flow across the bed is active throughout the water column, with the result that there is very little stratification. Nevertheless, there can be considerable lateral differences across the estuary in salinity and in mean flow velocity, as if the vertical circulation were turned on its side. This results in the currents on one side of the channel being flood-dominated while those on the other side are ebb-dominated, often


302


separated by a mid-channel bank (Figure 3). As a consequence, during the tide the water tends to flow preferentially landward up one side of the bank, cross the channel at high water, and ebb down the other side. The bed sediment also is driven to circulate around the bank in the same way. The above descriptive classification shows the general relationship between the structure of an estuary, the tidal range, and the river flow. Quantified classifications depend on producing numerical criteria that relate these variables. When the flow ratio the ratio of river flow per tidal cycle to the tidal prism volume – is 1 or greater, the estuary is highly stratified. When it is about 0.25, the estuary is partially mixed; and for less than 0.1 it is well mixed. Alternatively, the ratio of the river discharge velocity (R/A, where R is the river volume discharge, and A is the cross-sectional area) divided by the root-meansquare tidal velocity gives values of less than 102 for well-mixed conditions and greater than 101 for stratified conditions. There are many alternative proposals that depend on describing the processes controlling the tidally averaged vertical profiles of

5

10

15

20

salinity and of current velocity. One example uses a stratification parameter, which relates the surface to near-bed salinity difference (DS) normalized by the depth mean salinity (S0), and a circulation parameter that expresses the ratio of the surface mean flow (us) to the depth mean flow (uf) (Figure 2). Use of these classification schemes shows that estuaries plot as a series of points depending on position in the estuary, on river discharge, and on tidal range. They form part of a continuous sequence, rather than falling into distinct types, and an estuary can change in both space and time. Near the head of the estuary the river flow will be more important than nearer the mouth, and consequently the structure may be more stratified. Alternatively, if the tidal currents rise toward the head of the estuary, better-mixed conditions may prevail. Figure 4 shows the classification scheme based on the stratification–circulation parameters, together with indication of the direction in which an estuary would plot for different circumstances. Changes in river discharge of water force the estuarine circulation to respond. An increase will push the salt intrusion downstream and increase the stratification. Conversely, a decrease will allow the salt intrusion to creep further landward, and the stratification will decrease because the tidal motion will become relatively more important than the stabilizing effect of the fresh water input. Because of the drastic variation of tidal elevation and currents between spring and neap tides, well-mixed conditions

(A)

5

10

15

20

(B)

Depth

Salt

1.0

0.1

0.01

Well mixed

Mean velocity 0

Mean salinity

Stratification parameter ΔS / S0

10

1 LHS

RHS

RHS

Stagnant bottom layer wed ge

Partially mixed _ R _ X X+

10

LHS

Fiords

R+

102 103 104 Circulation parameter us /uf

105

(C)

Figure 3 Diagrammatic plan views of the horizontal distribution of salinity in a well-mixed estuary (PSU). (A) At mid flood tide. (B) At mid ebb tide. Contours show representative salinities; arrows show relative strengths of currents. (C) Vertical mean profiles for left-hand side and right-hand side of the estuary looking down estuary.

Figure 4 Quantified estuarine classification scheme based on stratification and circulation parameters. R and R þ show the direction of movement of the plot of an estuarine location with increase or decrease in river discharge; X and X þ with movement toward the estuarine head or mouth. Parameters are defined in Figure 2. From Dyer (1972) with permission. ^ John Wiley & Sons.



may occur at spring tides and partially mixed at neaps. This must occur by a relatively stronger mixing on the flood tide during the increasing tidal range, and by increased stabilization on the ebb during decreasing tidal range. The change between the two conditions may occur very rapidly at a critical tidal range as the effects of mixing rapidly break down the stratification. Additionally, a wind blowing landward along the estuary will tend to restrict the surface outflow, and may even reverse the sense of the vertical gravitational circulation. A down-estuary wind would enhance the stratification. Also, atmospheric pressure changes will cause a total outflow or inflow of water, as the mean water level responds. Thus large variations in the mixing and water circulation are likely on the timescale of a few days, the progression rate of atmospheric depressions.

Flushing Within the estuary there is a volume of fresh water that is continually being carried out to sea and replaced by new inputs of river water. At the mouth, not all of the brackish water discharged on the ebb tide returns on the flood, the proportion being very variable depending on the coastal circulation close to the mouth, and fresh water comprises a fraction of that lost. The flushing or residence time is the time taken to replace the existing fresh water in the estuary at a rate equal to the river discharge. From direct measurements of the salinity distribution it is possible to calculate the volume of accumulated fresh water and the flushing time for various river flows. The flushing time changes rapidly with discharge at low river discharge, but slowly at high river flow, being buffered to a certain extent by the consequent changes in stratification. Flushing times are generally of the order of days to tens of days, rising with the size of the estuary.

Turbidity Maximum A distinctive feature of partially and well-mixed estuaries is the turbidity maximum. This is a zone of high concentrations of fine suspended sediment, higher than in the river or lower down the estuary, that depends upon the estuarine circulation for its existence. The maximum is located near the head of the salt intrusion (Figure 5), and the maximum concentrations tend to rise with the tidal range, from of the order of 100 ppm to 10 000 ppm, though there are variations depending on the availability of sediment. Often the total amount of suspended sediment

303

5 10

15

20

(A)

60 80 100

40

(B)

Figure 5 (A) Diagrammatic longitudinal distribution of tidally averaged salinity and current velocities in a partially mixed estuary. Contours show salinities; arrows show relative current velocities. Dashed line shows level of no net motion. (B) Diagrammatic distribution of mean suspended sediment concentration (in ppm), showing the turbidity maximum.

exceeds several million tonnes. The location of the maximum coincides with a mud-reach of bed sediment, and with muddy intertidal areas. The continual movement of the turbidity maximum and the sorting of the sediment appears to create a gradient of sediment grain size along the estuary. Closer to the mouth, the bed is generally dominated by more sandy sediment. Because of the affinity of the fine particles for contaminants, prediction of the characteristics of the turbidity maximum is important. The position of the maximum changes with river discharge and with tidal range, but with a time lag behind the variation in the salt intrusion position. During the tide there are drastic changes in concentration. At high water the maximum is located well up the estuary and concentrations are relatively low because of settling. During the ebb the current reentrains the sediment from the bed, and advects it down the estuary. Settling again occurs at low water and there is further reentrainment and advection on the flood tide. Thus there is active cycling of sediment between the water column and the bed, and the whole mass becomes very well sorted in the process. The behavior of the sediment will depend on the settling velocity and the threshold for erosion. The settling velocity varies with concentration and with the intensity of the turbulence because flocculation of the particles occurs, forming large, loose but fastsettling floes. Settling causes high-concentration layers to build up on the bed, and at neap tides these may persist throughout the tide, forming layers often seen on echo-sounders as fluid mud. The turbidity maximum is maintained by a number of trapping mechanisms. Vertical gravitational circulation carries the fine particles and flocs downstream on the surface, and they settle toward the bed


304


because the turbulent mixing is reduced by the stratification. Near the bed the landward mean flow carries the suspended particles toward the head of the salt intrusion, together with new material coming in from the sea. There they meet other particles carried downstream by the river flow and are suspended by the high currents. Additionally, the asymmetry of the currents during the tide leads to tidal pumping of sediment, the flood-dominant currents ensuring that the sediment transport on the flood exceeds that on the ebb. The sediment is thus pumped landward until the asymmetry of the currents is reversed by the river flow – somewhere landward of the tip of the salt intrusion. The pumping process is affected by the settling of sediment to the bed and its reerosion, which introduces phase lags between the concentration and current velocity variations and enhances the asymmetry in the sediment transport rate during the tide. Measurements have shown that all of these processes are important, but tidal pumping is dominant in wellmixed estuaries. Both vertical gravitational circulation and tidal pumping are important in partially mixed estuaries. As a result the sediment in the turbidity maximum is derived from the coastal sea as much as from the rivers. In microtidal salt wedge estuaries the surface fresh water carries sediment from the river seaward throughout the estuary, discharging into the sea at high river flow. The clearer water in the underlying salt wedge is often revealed in the wakes of ships.

Estuarine Modeling The water flows transport salt, thus affecting the density distribution, and the densities affect the longitudinal pressure gradients that drive the water flow. Turbulent eddies produced by the flows cause exchanges of momentum as well as of salt, and produce frictional forces that help to resist the flow. Estuarine models attempt to predict the salt distribution, the water circulation and the sediment transport through application of the fundamental equations for mass, momentum, and sediment continuity, which formalize the above interactions. In principle, it is possible to measure directly most of the terms in these equations, apart from those that involve the turbulent exchanges. In practice, approximations are required to solve the equations. Either tidally averaged or within-tide conditions can be modeled. For sediment, the definition of the settling velocity and the threshold are required. One-dimensional (1D) models assume that the estuarine characteristics are constant across the

cross-section, and that the estuary is of straight prismatic shape. For tidally average conditions the assumption is that the seaward advection of salt on the mean flow is balanced by a landward dispersion of salt at a rate determined by the horizontal salinity gradient and a longitudinal eddy dispersion coefficient. This coefficient obviously incorporates the effects of turbulent mixing, the tidal oscillation and any actual nonuniformity of conditions. Two dimensional (2D) models can either assume that conditions are constant with depth but vary across the estuary (2DH), or are constant across the estuary but vary with depth (2DV). The eddy dispersion and eddy viscosity coefficients will be different for the two situations. Using the 2DH model one would not be able to explore the effects of vertical gravitational circulation, whereas with the 2DV model the effects of the shallow water sides of the channel would be missing. In many cases a further approximation may be made by assuming the water is of constant density. Three-dimensional models are obviously the ideal as they represent fully the vertical and horizontal profiles of velocity, density, and suspended sediment concentration. The only unknown terms are those relating to the turbulent mixing, the diffusion terms. Although computationally 3D models are becoming less costly to run, it is still difficult to obtain sufficient data to calibrate and validate them. Once realistic flow models have been constructed, they can be used to explore the transport of sediment and estuarine water quality. However, the limitations on the field data restricts their use for prediction. For instance, the majority of the sediment travels very near to the bed, and it is difficult to measure the concentrations of layers only a few centimeters thick. An active topic for research and modeling at present is the prediction of morphological change in estuaries resulting from sea level rise. This involves integrating the sediment transport over the tide, for many months or years, in which case the accumulated errors can become overriding. These results can then be compared with the simple empirical relationships stemming from considerations of estuarine equilibrium.

See also Breaking Waves and Near-Surface Turbulence. Coastal Circulation Models. Elemental Distribution: Overview. Fiord Circulation. Fiordic Ecosystems. General Circulation Models. Intrusions. NonRotating Gravity Currents. Rotating Gravity Currents. Tides. Upper Ocean Mixing Processes.



Further Reading Dyer KR (1986) Coastal and Estuarine Sediment Dynamics. Chichester: Wiley. Dyer KR (1997) Estuaries: A Physical Introduction, 2nd edn. Chichester: Wiley. Kjerfve BJ (1988) Hydrodynamics of Estuaries, vol. I: Estuarine Oceanography; vol. II: Hydrodynamics. Boca Raton, FL: CRC Press.

305

Lewis R (1997) Dispersion in Estuaries and Coastal Waters. Chichester: Wiley. Officer CB (1976) Physical Oceanography of Estuaries (and Associated Coastal Waters). New York: Wiley. Perillo GME (1995) Geomorphology and Sedimentology of Estuaries. Amsterdam: Elsevier.


EUTROPHICATION V. N. de Jonge, Department of Marine Biology, Groningen University, Haren, The Netherlands M. Elliott, Institute of Estuarine and Coastal Studies, University of Hull, Hull, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 852–870, & 2001, Elsevier Ltd.

Introduction Eutrophication is the enrichment of the environment with nutrients and the concomitant production of undesirable effects, while the presence of excess nutrients per se is merely regarded as hypernutrification. In more detail, eutrophication is the process of nutrient enrichment (usually by nitrogen and phosphorus) in aquatic ecosystems such that the productivity of the system ceases to be limited by the availability of nutrients. It occurs naturally over geological time, but may be accelerated by human activities (e.g. sewage disposal or land drainage). (Oxford English Dictionary)

Anthropogenic nutrient enrichment is important when naturally productive estuarine and coastal systems receive nutrients from ‘point sources’, e.g. as outfall discharges of industrial plants and sewage treatment works, or human-influenced ‘diffuse

% area: % production:

LAND

sources’, such as runoff from an agricultural catchment. Whereas point source discharges are relatively easy to control, with an appropriate technology, diffuse and atmospheric sources are more difficult and require a change in agricultural and technical practice. Coastal and estuarine areas with tide-associated accumulation mechanisms for seaborne suspended matter are organically productive by nature and represent some of the world’s most productive environments. This is the result of the freshwater outflow, and biogeochemical cycling within the estuarine systems and adjacent shallow coastal seas. About 28% of the total global primary production takes place here, while the surface area of these systems covers only 8% of the Earth’s surface (Figure 1) and as such the effects of eutrophication are most manifest in the coastal zone, including estuaries, areas which are the focus of this chapter. As indicated below, there is generally a good qualitative understanding of the processes operating, but the quantitative influence on the ecological processes and the changes in community structure are still not well understood. Within the available field studies attention has been focused on long data series, because time-series with a length of less than 10 years have to be considered as too narrow a window of time relative to natural meteorological and climatic fluctuations influencing the ecosystem.

8 26

27 41

COASTAL ZONE

65 33

OCEAN

Figure 1 Indication of importance to primary production of different typesof area (land, ocean, and shallow coastal seas and its fringes). (Redrawn after LOICZ, 1993. Report no. 25, Science Plan, Stockholm.)

306


EUTROPHICATION

Nutrient inputs are required for the natural functioning of aquatic systems. Eutrophication merely indicates that the system cannot cope with the available inputs. This chapter focuses on the causes and mechanisms of eutrophication as well as the consequences. It gives examples varying from eutrophication caused by freshwater inflow to that caused mainlyby atmospheric inputs and nutrient import from the sea instead of land and atmosphere. It will be shown that increased organic enrichment may lead to dystrophication, the modification of bacterial activity leading ultimately to anoxia. Despite this, certain areas with low hydrodynamic energy conditions, such as lagoons, part of the estuaries and enclosed seas, can be considered as naturally organically enriched and thus require little additional material to make them eutrophic. In contrast, there are also naturally oligotrophic areas which drain poor upland areas and receive little organic matter. The symptoms of eutrophication as a response to nutrient enrichment differ greatly, due mainly to differences in the physical characteristics of the different systems receiving this ‘excess’ organic matter and nutrients. For example, the mixing state (stratification or not) of the receiving body of water and theresidence time (or flushing time) of the fresh water and its nutrients in the system determine the intensity of a particular symptom and thereby the sensitivity of systems to eutrophication events or symptoms. For example, in the UK, the susceptibility of waters to enrichment and adverse effects caused by nutrients is interpreted according to their ability as high natural dispersing areas (HNDA). In general, this reflects the waters’ assimilative capacity, i.e. capacity to dilute, degrade, andassimilate nutrients without adverse consequences.

307

the early 1950s, due to the introduction of artificial fertilizers and detergents. Available data for Narragansett Bay (USA) suggest that the system presumably was nitrogen limited in thepast and that the total dissolved inorganic nitrogen (DIN) input has increased five-fold, while that of dissolved inorganic phosphorus (DIP) has increased two-fold (Table 1). During that period the nutrient input from thesea is assumed to have been more important than the supply from the drainage basin, a feature that in the early 1950s has also been postulated for European waters. Natural background concentrations for someDutch and German rivers as well as the Wadden Sea are given in Table 1. The values represent the situation before the introduction of the artificial fertilizers and detergents. The 4- to 50-fold increase in the values in fresh waters and coastal waters is clear.

Structuring Elements and Processes In addition to the inputs, the mechanisms and processes which influence the fate and effects of excess nutrient inputs will be considered. In coastal systems, eutrophication will influence not onlythe nutrientrelated processes of the system but will also affect structural elements of the ecosystem (Figure 2).

Structuring Elements

The physical and chemical characteristics mainly create the basic habitat conditions and niches of the marine system to be colonized with organisms. These conditions also determine colonization rate, which is dependent on the organisms’ tolerances to environmental variables.

Historical Background Several attempts have been made to determine either the nutrient loads to coastal zones during the ‘predevelopment period’ (i.e. before widespread human development) or to determine the ‘natural background’ concentrations of nutrients. These assessments are important, as they may improve our understanding of the way eutrophication in the past may havechanged and in the future will change aquatic systems. The term ‘natural’ is regarded here as those levels present before the large-scale production and use of artificial fertilizers and detergents started. It also covers the period just prior to the start of chemical monitoring of the aquatic environment. For example, the greatest increase innutrient loads to the Dutch coast and the Wadden Sea occurred after

Processes

Many nutrient transformation processes occur in estuaries, as they have the appropriate conditions. Estuaries can be considered as reactor vessels with a continuous inflow of components from the sea, the river, and the atmosphere and an outflow of compounds to the sea, the atmosphere, and the bottom sediments after undergoing certain transformations within the estuary (Figure 3). The important process elements are import, transformation, retention, and export of substances related to organic carbon and nutrients. Part of the transformation processes is illustrated in Figure 4. All these processes dictate that estuaries should be considered as both sources and sinks of organic matter and nutrients.


308

EUTROPHICATION

Table 1 Recent and ‘prehistoric’ loadings of some systemsby nitrogen and phosphorus and the resultant annual primary production of these systems System: period/year

P influx (mmol m 2 a 1)

N influx (mmol m 2 a 1)

River Rhine and River Ems ‘background’ situation English Channel ‘background’ situation North Sea ‘background’ situation Dutch western Wadden Sea ‘background’ situation

Mean annual nitrogen concentration in system (mmol l 1)

Mean annual phosphorus concentration in system (mmol l 1)

45 7 25 (tN)

1.8 7 0.8 (tP)

5.5 7 0.5 (winter NO3)

0.45 7 0.05 (winter DIP)

9.1 7 3.1 (NO3) (near coast)

0.57 7 0.13 (DIP) (near coast)

13 7 6 (tN) c. 4 (DIN) (for salinity gradient) 10–45 (tN)

0.8 7 0.3 (tP) c. 0.3 (DIP) (for salinity gradient) 0.7–1.8 (tP)

Annual primary production (g C m 2 a 1)

o50

Ems estuary

(rivers)

(rivers)

‘background’ situation early 1980s early 1990s Baltic Sea ‘background’ situation (c. 1900) early 1980s

315 (tN)

16 (tP)

3850 (tN) 3850 (tN)

90 (tP) 50 (tP)

57 (tN) (rivers þ AD þ fix) 230 (tN) (rivers þ AD þ fix)

0.8 (tP) (rivers þ AD) 6.7 (tP) (rivers þ AD)

80–105 135

18–76 (DIN) (rivers þ AD) 270–330 (DIN)( þ sea input)

130

1445 (DIN) (rivers þ AD) 1725 (DIN)( þ sea input)

B1 (DIP) (rivers þ AD) 61 (DIP)( þ sea input) 73 (DIP) (rivers þ AD) 140 (DIP)( þ sea input)

— 1040 (tN) (rivers þ AD) 290–2140 (DIN) 1430 (tN)

— 70 (tP) (rivers þ AD) 30 (DIP) 40 (tP)

Narragansett Bay ‘prehistoric’ situation

‘recent’ (1990s)

Long Island Sound 1952 early 1980s Chesapeake Bay mid-1980s

290

c. 200 300 400–600

tN, total nitrogen; tP, total phosphorus; DIN, dissolved inorganic deposition nitrogen; DIP, dissolved inorganic phosphorus; AD, atmospheric deposition; fix, nitrogen fixation.

Output

Several processes contribute significantly to the prevention of eutrophication symptoms. Active bacterial removal of nitrogen may occur under favorable conditions due to the conversion of nitrogen compounds into nitrogen gas. These conditions are the spatial change from anoxic and hypoxic and oxygenated conditions. Phosphorus may be removed

from the system due to either transport to the open sea or the permanent burial of apatites, for example. There is a wide variation in bacterial processes enhancing the transformation of nutrients. The intensity of the several conversion processes is partly related to the differences in the dimensions of these systems and factors such as tidal range (energy) and


EUTROPHICATION

309

Biological factors Ecosystem structure – niche creation Geological factors – substratum

Biological factors Ecosystem structure – species composition – species diversity – species properties – K or r-strategists

Habitat formation

Physical factors – tide – river discharge – salinity – turbidity – temperature

Biological factors Processes – net autotrophy – net heterotrophy

Chemical factors – availability of N & P

Biological factors Feedback mechanisms varying geo-chemical conditions – Eh – pH

Figure 2 Operating forcing variables in the development of estuarine and marine biological communities.

ATMOSPHERIC DEPOSITION

EMISSION

dissolved Fe, dissolved Al, dissolved C organic, DIP

EXPORT to sea (dissolved) Estuary C B A Transformation IMPORT from sea (particulate)

river IMPORT (part. + diss.) RETENTION by burial TIDAL FRESH RIVER

Figure 3 Box model showing fluxes, gradients, and processes in the freshwater tidal river, the estuary, and the coastal zone of the sea; part., particulate; diss., dissolved.

freshwater inflow (flushing time of fresh water, residence time of fresh or sea water and turnover time of the basin water) and related important determinants such as turbidity. The most important factors in the expression of eutrophication are: flushing time (ft), the turbidity

(gradient) expressed as light extinction coefficient (kz), and the input and concentration gradient of the nutrients N, P, and Si. The combination of mainly these three factors determines whether an estuarine system has a low or high risk of producing eutrophication symptoms (Figure 5).


310

EUTROPHICATION

NO2– NO3–

N2

NH4+ PO43 –

Exchange of nutrients

N 2O

Alga NO2–

WATER PHASE SEDIMENT Oxygenated photic zone NO3

Denitrification

PO43 –

NH4+

–

NH4+

P

PO43 –

Feadsorbed

P

PO43 –

Porganic

Anoxic

Feadsorbed

Norganic

NO2–

Oxygenated

PO43 –

Figure 4 Schematic representation of conversion processes of nitrogen and phosphorus. Dashed lines represent assimilation, full lines represent biotic conversion (mainly microbial), and dotted lines represent geochemical equilibria. (Modified after Wiltshire, 1992 and van Beusekom & de Jonge, 1998 and references therein.)

Nutrient input

High High risk

Low

High

Water clearness Low

Low risk Low

Flushing time

High

Figure 5 A 3-D classification scheme of eutrophication risk of estuaries based on flushing time, turbidity, and nutrient input.


EUTROPHICATION

The longer the flushing time then the more vulnerable the system is to nutrient enrichment, as the primary producers have a greater period available to utilize the excess nutrients. If the flushing time is shorter than the mean growth rate of the algae (areas with high dispersion capacity), flushing of the population will occur and thus prevent the symptoms within the system, although transport to the open sea will increase. Hence, if algal blooming does not occur within the estuary it may develop in the lowest reaches of the system or even just outside the system in the sea.

Case Studies Indicating Trends and Symptoms In determining any response, it is necessary to describe the natural situation and its spatial and temporal variability, the change from that natural system, and the significance and cause of that change. As described above, it may be difficult to assess anthropogenic nutrient enrichment against a

311

background of the natural variability in nutrient influxes and turnover and its impact on the productivity of the ecosystem under consideration. The greatest problem is to make a clear distinction between the two and to assess the contribution of natural variation and of human activitiesto the nutrient enrichment and its symptoms in any system. Despite this, there are several case studies which illustrate the main features, as describedbelow. Forcing Variable 1: Riverine Inputs and Concentrations of Nutrients

There are many scattered data available on input values and concentrations of nutrients, but there are much less consistent long-term data series. Comparisons of historical and present data for the North Sea and Wadden Sea indicate that large-scale variability in inputs (Figure 6; inflowing Atlantic water) and large increases in inputs (Figure 7; inputs from rivers) resulted in an increased productivity. An analysis of the river, estuarine and coastal dynamics and an understanding of the processes, especially in

NO3− + NO2− (mmol m−3) 30 25 20 15 10 5 0 1930

1950

1970

1990

1950

1970

1990

PO43− (mmol m−3) 1.6 1.4 1.2 1.0 0.8 0.6 0.4 0.2 0 1930

Figure 6 Winter concentrations of DIN and DIP in the central part of the Strait of Dover showing a two-fold increase in DIN and a three-fold increase in DIP. (Reproduced with permission from de Jonge, 1997).


312

EUTROPHICATION

Mean load of total P (mol s−1) 70 Rhine at Lobith 60

50

40

30

20

10

0 1950

1955

1960

1965

1970

1975

1980

1985

1990

1995

1970

1975

1980

1985

1990

1995

Mean load of total N (mol s−1) 1600 Rhine at Lobith

1400 1200 1000 800 600 400 200 0 1950

1955

1960

1965

Figure 7 Development of the loads in total phosphorus and total nitrogen as measured on the border between Germany and The Netherlands. (Data from Rijkswaterstaat).

the near-coastal dynamics, showed that there may be a time delay inherent in the system and that the coast may respond later than the estuary. It is of note that detailed and systematic monitoring was required to detect these trends. Consequently, international agreements (the Oslo and Paris Conventions and European Commission Directives) were designed to control these trends. The measures were effective and

thus the remediation may be a model for systems elsewhere (cf. Figure 2). The changes reported include an increase of surface algal blooms (Figure 8), a major sudden change in plankton composition from diatoms to small flagellates, and an increase in the area with elevated nitrogen and phosphorus concentrations (Figure 9) and oxygen deficiency (Figure 10). There are well-defined


EUTROPHICATION

313

Number of blooms 70 60 50 40 30 20 10 0 1979

1983

1981

1985

1987

1989

1985

1987

1989

1991

1993

1995

Surface area (km2) 12 000 10 000 8000 6000 4000 2000 0 1979

1981

1983

1991

1993

1995

Figure 8 Surface algal blooms observed during Dutch airborne surveys over the period 1979–1995 (after Zevenboom 1993, 1998). (A) Frequency per annum; (B) surface area in km2 per annum.

positive relationships between riverborne nutrients and the long-term variation in primary and secondary production in the western Dutch Wadden Sea (Figure 11). However, the processes and responses are complicated by inputs from the North Sea, including the inflowing Atlantic water (Figures 6 and 12) and meteorological conditions. This is important as it emphasizes the need to consider all aspects when reduction measures have been undertaken. Forcing Variable 2: Size of Receiving Area

The Baltic Sea is typical of many semi-enclosed seas with nutrient loadings. Although the four- to eightfold increased nutrient loads only changed the primary production by 30–70%, the permanent anoxic

layer of the Baltic increased from 19 000 to near 80 000 km2 (nearly the entire hypolimnion) as the result of the eutrophication (Figure 13). This produced a dramatic structural decline in population size of two important prey species (the large isopod Saduria entomon and the snake blenny Lumpenus lampetraeformis), which in turn greatly influenced the local cod populations. The increased density of algal mats reduced the development of herring eggs, possibly by exudate production and the large hypoxic areas adversely affected the development of cod eggs. Thus nutrient enrichment in the Baltic greatly damaged some essential parts of the food web and the ecosystem structure. In additon, more localized eutrophication of the inner Baltic occurs due to fish farming (Figure 14),


314

EUTROPHICATION

Norway

North Sea DK

Germany UK The Netherlands

8 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol NO3 l −1; > 0.8 µmol P l−1 Elevated nutrient levels Temporarily elevated NO3 concentration >10 µmol N l−1; > 8 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol NO3 l −1; > 0.8 µmol P l−1 > 15 µmol N l −1 > 0.8 µmol P l−1

Figure 9 Elevated nutrient levels in North Sea waters (data from OSPAR 1992). The background levels have been agreed to be 10 mmol l1 DIN and 0.6 mmol l1 DIP (from Zevenboom 1988, 1993).

which accounts for over 35% and 55% of the total local nitrogen and phosphorus inputs respectively. It was concluded that phosphorus was the most important determinant, e.g. leading to a strongly reduced N/P ratio. This nutrient enrichment produced an increase in the primary production and an increase in turbidity which negatively affected the macrophyte populations and stimulated the blooming of cyanobacteria. The loss of five benthic crustaceans has been observed, while a gain of four was reported, of which two were polychaetes (Polydora redeki, Marenzelleria viridis) new to the area. Furthermore, the macrobenthos showed a structural change from suspension feeders to deposit feeders. Finally, the extensive bloom of the microalga Chrysochromulina in the late 1980s in the outer Baltic

apparently developed in response to nutrient buildup on the eastern North Sea and contributed to the hypoxia and eutrophic symptoms. Forcing Variable 3: Peak Loadings and Changing Ratios in Nutrient Fractions

In Chesapeake Bay (USA; north Atlantic west coast) (Figure 15), the total phosphorus concentrations decreased with time, but nitrogen had maximum concentrations in the mid-1980s and the DIN concentrations doubled in the oligohaline zone of the bay. Significant increases were detected in surface chlorophyll-a data between 1950 and 1970 in all the regions of the bay and there has been a significant long-term increase in the DIN/DIP ratios since the


EUTROPHICATION

315

Norway

North Sea DK

Germany UK The Netherlands

-

O2 deficiency Wadden Sea sediments (since 1988) O2 deficiency sediments (1989) < 2 mg O2 l −1 (1980–89) < 2 mg O2 l −1 (1981–90) 4–8 mg O2 l −1 (1989–90) 4–8 mg O2 l −1 (1980–90) 6–7 mg O2 l −1 (1979–90)

Figure 10 North Sea areas with oxygen deficiency. (Data after OSPAR, 1992; Zevenboom, 1993).

1960s in much of the bay. This ratio was generally above the Redfield ratio of 16 (necessary for optimal plant growth) in all regions in winter and spring, but in summer and autumn the values were below the Redfield ratio in the main part of the bay, suggesting N limitation. Potentially limiting concentrations of reactive silicate and DIP often occurred in the mesohaline to polyhaline bay. Forcing Variable 4: River Basin Alterations

Alterations in the lower part of the Rhine river basin have led to a decrease in the flushing time of the system and concurrent relative increase in the discharge of nutrient loads to the Dutch Wadden Sea. Eutrophication in Florida Bay (Figure 16) possibly produced a large-scale seagrass die off (4000 ha of

Thalassia testudinum and Halodule wrightii disappeared between 1987 and 1988), followed by increased phytoplankton abundance, sponge mortality, and a perceived decline in fisheries. This very large change in the health of the system followed major engineering works to the Everglades area and reflected changes in the nutrient concentrations, the nutrient pool, the chlorophyll levels, and turbidity. The preliminary nutrient budget for the bay assumes a large oceanic and atmospheric input of N and P to the bay, although the denitrification rates are unknown. The cause(s) of the seagrass mortality in 1987 is still unknown. Although not mentioned, synergistic effects (where eutrophication effects are exacerbated by other pollutants) are also expected in urbanized and developed areas. Furthermore, the change in primary producers in this area reflected


316

EUTROPHICATION

Primary production (g C m−2 a−1) 600 Phytoplankton Marsdiep 500

400

300

200

100

* 0 1950

1960

1970

1990

1980

Figure 11 Time-series of available values on primary production in the western Dutch Wadden Sea, as reviewed by de Jonge et al. 1996 (with permission).

(μmol l−1) 10.0 Rhine 1986 DIP 7.5

5.0 Rhine 1990 2.5

2.5

Lake Ijssel 1986 Lake Ijssel 1990

0

0 0

10

20

Salinity English

30 Channel

Figure 12 Mixing diagram of DIP for the years 1986 and 1990 when nutrient concentrations in river water declined after having increased for decades. The strong and structural (cf. also main river Rhine values in Figure 13) decrease in DIP values over a short time period is remarkable. Also the conclusion that in 1990 the DIP values of fresh water in Lake Ijsselmeer were lower than the values in theStrait of Dover/English Channel (which has consequences for the primary production potential of coastal waters in the southern Bight of the North Sea and coastal policy plans and management plans is striking. (Reproduced with permission from de Jonge, 1997.)


EUTROPHICATION

317

Phosphate (μmol l–1)

Nitrate (μmol l–1) 10 8 6 4

5 1

2 0 1960

0 1960

1980

1970

1980

6 1 4 2

0 1960 1970

4

0 1960 1970

1980

1980 3

10 8 6 4 2 2

2

0 1960 1970

1980 1 1 0 1950 1960

1970

1980

4 1 2 0 1960

1970

1980

0 1960

1970

1980

Figure 13 Map of the Baltic Sea with hypoxic and anoxic areas in (1) Arkona Basin, (2) the Gotland Sea, and (3) Gulf of Finland. Further trends in nitrate (left panels) and phosphate (right panels) in surface waters in winter (gray) and at 100 m depth (black). (Reproduced with permission from Elmgren, 1989.)

vascular plants operating as k-strategists under stress and replaced by more opportunistic algae like phytoplankton species. Forcing Variable 5: Stratification

The Peel-Harvey estuarine system (South Pacific west coast of Australia) receives a high nutrient loading but has a phosphorus release during stratification-

induced anoxia from the bottom sediments. This is after a clear loading of the estuary and the subsequent development of dense populations of microphytobenthos which is responsible for the nutrient storage. This release (Figure 17) contributed to the changes in macroalgal community structure and increased turbidity due to algal blooms. Remedial measures are now in place to reduce inputs and remove the adverse symptoms.


318

EUTROPHICATION

Kumlinge, Åland islands (winter) 30

350

25

300 20

250 200

15

150 100 50

tot-N

r 2 = 0.686; p < 0.01

10

tot-P

r 2 = 0.728; p < 0.01

5

0

tot-P (1–10 m)

tot-N (1–10 m)

400

0 65 67

69

71

73

75

77

79 81

83

85

87

89

91

Year

93

1993

1991

1989

1987

1985

1983

1981

1979

1977

1975

1973

1971

1969

1967

1965

Rajakari, the Airisto Sound

0

secchi (m)

0.5 1 1.5 2 2.5 3 3.5 r 2 = 0.313; p < 0.01

4 4.5 400 outer r 2 = 0.759; p < 0.01

Primary production capacity

350

central r 2 = 0.581; p < 0.05

300 250 200 150 100 50 0 76

78

80

82

84

86

88

90

Year

Figure 14 Development of nutrient levels in the open outer archipelago region (Baltic part of Finland) over the period 1963–93, the increase in turbidity over the period 1965–93, and the consequent increase in primary production capacity over the period 1976–90. (Reproduced with permission from Bonsdorff et al., 1997.)

Forcing Variable 6: Hydrographic Regime (Residence Time and Turbity)

The creation of red tides (noxious, toxic, and nuisance microalgal blooms) in Tolo Harbor (Hong Kong) resulted from large urban nutrient inputs, a water

residence time of 16–42 days, and a low turbidity which led to the phytoplankton producing dense populations. Diatoms decreased in abundance from 80–90% to 53% in 1982–85, dinoflagellates increased concurrently with red tides (Figure 18) and


EUTROPHICATION

N loading to bay

NoX

DON

NH4

PN

2.2

0.6

0.4

1.8 1.6 1.4

0.2

1.2 0.14

1988

1987

1986

1985

1984

1983

1982

1981

1980

1979

1978

0.0

TPP DOP PO4

0.4

mg P I−1

P loading to bay

TP concentration

0.12

Years 0.5 Phosphorus loading, kg P × 105 d−1

TP concentration

2.0 mg N I−1

Nitrogen loading, kg N × 106 d−1

0.8

319

0.10 0.08 0.06 0.04 0.02

0.3

0.2

0.1

1988

1987

1986

1985

1984

1983

1982

1981

1980

1979

1978

0.0

Years Figure 15 Record of monthly averaged nitrogen and phosphorus inputs to the mainstem of Chesapeake Bay and the monthly average values of total nitrogen and total phosphorus measured at the fall-line of the Susquehanna River (original data from Summers 1989). (Reproduced with permission from Boynton et al., 1995.)

chlorophyll-a levels also increased significantly. Oxygen depletion occurred due to nutrient loadings and the local development of phytoplankton in combination with reduced water exchange (long flushing time), features associated with a low-energy eutrophic environment. In addition, the area is degraded after heavy pollution by heavy metals and organic contaminants. In other systems (e.g. Southampton Water, UK) regular blooms of the nuisance ciliate with a symbiotic red alga, Mesodinium rubrum, is the result of increased nutrients, high organic matter, relatively long residence time, and turbid waters.

Conceptual Model of Effects There have been few studies assessing all possible effects of nutrient enrichment but it is possible through the many different case studies to create a

conceptual model of the main effects. Figure 19 gives a descriptive overview on functional groups of plants and animals and at differing levels of biological organization, thus mainly at the process level.

‘Hot Spots’ and Remedial Measures With regard to eutrophication, ‘hotspots’ may be those being hypernutrified, such as estuaries (e.g.the Ythan, Scotland) or those areas showing regular symptoms ofeutrophication, e.g. the Baltic Sea. Other good examples are the near absence of beaver dams in the USA today, and the absence of large natural wetlands as aresult of reclamation in many low-lying countries. In the past these natural obstacles as beaver dams and large wetlands favored the retention of nutrients resulting in lower more ‘near’ natural loads of coastal systems. It is clear that


320

EUTROPHICATION

Figure 16 Map of Florida Bay with three zones of similar influence as a result of a cluster analysis on mean and SD of five principal component scores. Stations are labeled as eastern bay ( þ ), central bay (’), and western bay ( ). (Reproduced with permission from Boyer et al., 1999.)

restoration of river systems or the rehabilitation of the integrity of entire river systems in combination with the application of best possible techniques is the best remedial measure to implement, coupled with river basin and catchment management. In general ‘hot spots’ are allclose to intensive land use (agriculture and urbanized areas), with poor waste water treatment and no removal of P and N. Increasing development is usually accompanied by greater waste treatment, for example, European Directives require better treatment depending on the local population and theability of receiving waters to assimilate waste. However, it is axiomatic that sewage treatment removes organic matter but, unless nutrient stripping is installed, which is expensive, it may fail to remove, or hardly remove nutrients. Similarly, the creation of nitrate vulnerable areas requiring fertilizer control, as within the EU Nitrates Directive, will

reduce inputs. However, the fact that ground water may retain nutrients for many years, even decades in the case of aquifers, will dictate that the results of remediation will not be apparent for a while. Areas requiring attention include populated regions, agricultural lands, and low-energy areas (Baltic Sea with A˚land Islands, German Bight in the North Sea, Long Island Sound,Chesapeake Bay), i.e. mainly the large estuarine systems as well as developing countries with no or hardly any wastewater treatment. Anthropogenic eutrophication must be addressed, especially further improvement of wastewater treatment and technical processes to reduce the emissions of nutrients and related (NOx) compounds to the atmosphere. Despite increasing knowledge, most countries show the same history when focusing on eutrophication. The fact that the information given above suggests a


EUTROPHICATION

321

50 100 45

40

80

30

60

25

40

Salinity (ppm)

Phosphate release (mg m−2 d −1)

35

20 15

10

20

5 0

J

O

J

A

J

O

J

A

J

O

1988

1989

J

A

0

1987

Figure 17 Release of phosphorus from the sediments of Harvey estuary. The data are release rates, recorded under standard conditions in the laboratory, for cores removed from the same site in the estuary at different times of year. The background line shows the estuarine salinity at the time. (Reproduced with permission from McComb, 1995.)

Number of events 45 40 35 30 25 20 15 10 5 0

1977

1979

1981

1983

1985

1987

1989

Figure 18 Red tides and associated fish kills in Tolo Harbor. (Reproduced with permission from Hodgkiss and Yim, 1995.)

reduction in the emission of nutrientsshould be interpreted with caution, because differences in nutrient ratios in combination with changes in concentrations may lead to the development of undesirable microand macro-algae. For example, Sweden’s reduction policy, which focused on phosphorus, failed as phosphorus became depleted along the coasts but not in the central part of the Baltic Sea where it was supplied in

excess from anoxic deep water – thus maintaining the near-surface algal blooms. Given the action plans adopted by developed nations to further reduce nutrient loads, it can be argued that in the near future, eutrophication will be caused by sea water that has been enriched with nutrients for decades instead of fresh water. This is due to the expectation that the present nutrient policy on ‘diffuse sources’ and the


LOW

(c) 2011 Elsevier Inc. All Rights Reserved. Nuisance macroalgal blooms

Noxious/nuisance microalgal blooms

Hypernutrification

LOW

HIGH

LONG

Toxic microalgal blooms

Disrupted intertidal bird feeding areas

Anoxia under algal mats

Decreased light penetration

High organic input to sediment

Reduced amenity

PSP, ASP, DSP

Uptake by filter feeders

Reduced palatibility of infauna

Reduced phytoplankton production

Sediment anoxia

Water anoxia

Vertebrate death

Opportunistic fauna: high abundance, moderate biomass, low diversity

Toxic response

Increased predation pressure in other areas

Reduced carrying capacity

Disrupted fish feeding

Sulphide and methane production

Changed fish community

Reduction in commercial fisheries

EXPORT TO SEA/OCEAN

Water quality barrier for fish movement

CONSEQUENCES

Figure 19 Overview of potential effects of nutrient enrichment in combination with turbidity conditions and residence time of water in an estuarine area.

Natural light condition

LIGHT

Levels and ratios of N&P

POOR

SYMPTOMS

BIOGEOCHEMICAL CYCLING GOOD

SHORT

RESIDENCE TIME

N P Si

HIGH

RIVER INFLUX

CONDITIONS

322 EUTROPHICATION

EUTROPHICATION

increasing application of modern, sophisticated wastewater treatment plants will further diminish the freshwater loads. However, the atmospheric deposition of nitrogen as well as phosphorus (in dust) will become increasingly important due to many nutrient sources resulting from land use (burning of fossil carbon, fields, and forests). The process of nitrogen fixation of increasing future importance as a mechanism during low nutrient conditions to compensate for the remedial measures taken by the different governments. This expectation means a well-balanced reduction in nutrient loads to prevent noxious blooms. It also means continuing to pay attention to eutrophication in all its aspects. At the moment nitrogen fixation is probably a small N-source as is the case in most nutrient-rich estuarine systems. However, some species have developed the ability to cope with very low nitrogen concentrations under conditions where just enough is provided by nitrogen fixation. Further global reduction in nitrogen emissions is required to protectthe environment. It is possible that the problem due to N fixation will be apparent when reduction in phosphorus loads have been taken as far as possible. The most important ‘hot spot’ on this planet is the rapidly growing world population. The big question and challenge is how to offer every individual ‘sustainable’ living conditions while at the same time maintaining the integrity of our aquatic systems. This marked increase in population size is the main cause of the most common and most severe environmental problem of today and tomorrow.

See also Carbon Cycle. Coastal Circulation Models. Marine Silica Cycle. Nitrogen Cycle. Phosphorus Cycle. Phytoplankton Blooms. Primary Production Distribution. Primary Production Processes. Regional and Shelf Sea Models. River Inputs. Tides.

Further Reading van Beusekom JEE and de Jonge VN (1998) Retention of phosphorus and nitrogen in the Ems estuary. Estuaries 21: 527--539. Boyer JN, Fourqurean JW, and Jones RD (1999) Seasonal and long-term trends in the water quality of Florida Bay (1989–1997). Estuaries 22: 417--430.

323

Boynton WR, Garber JH, Summers R, and Kemp WM (1995) Inputs, transformations, and transport of nitrogen and phosphorus in Chesapeake Bay and selected tributaries. Estuaries 18: 285--314. Daan N and Richardson K (eds.) (1996) Changes in the North Sea ecosystem and their causes: A˚rhus 1975 revisited. ICES Journal of Marine Science 53: 879--1226. de Jonge VN (1997) High remaining productivity in the Dutch western Wadden Sea despite decreasing nutrient inputs from riverine sources. Marine Pollution Bulletin 34: 427--436. de Jonge VN, Bakker JF, and van Stralen M (1996) Recent changes in the contributions of river Rhine and North Sea to the eutrophication of the western Dutch Wadden Sea. Netherlands Journal of Aquatic Ecology 30: 27--39. de Jonge VN and Postma H (1974) Phosphorus compounds in the Dutch Wadden Sea. Netherlands Journal of Sea Research 8: 139--153. Elmgren R (1989) Man’s impact on the ecosystem of the Baltic Sea: energy flows today and at the turn of the century. Ambio 18: 326--332. Hodgkiss IJ and Yim WW-S (1995) A case study of Tolo Harbour, Hong Kong. In: McComb AJ (ed.) Eutrophic Shallow Estuaries and Lagoons, CRC-Series, pp. 41--57. Boca Raton, FL: CRC Press. Libes SM (1992) An Introduction to Marine Biogeochemistry. New York: John Wiley Sons. McComb AJ (ed.) (1995) Eutrophic Shallow Estuaries and Lagoons. CRC-Series. Boca Raton, FL: CRC Press. Olausson E and Cato I (eds.) (1980) Chemistry and Biogeochemistry of Estuaries. New York: John Wiley & Sons. Salomons W, Bayne BL, Duursma EK, and Fo¨rstner U (eds.) (1998) Pollution of the North Sea: An Assessment. Berlin: Springer-Verlag. Stumm W (1992) Chemistry of the Solid–Water Interface. Processes at the Mineral–Water and Particle–Water Interface in Natural Systems. New York: John Wiley Sons. Summers RM (1989) Point and Non-point Source Nitrogen and Phosphorus Loading to the Northern Chesapeake Bay. Maryland Department of the Environment, Water Management Administration, Chesapeake Bay Special Projects Program, Baltimore, MD. Wiltshire KH (1992) The influence of microphytobenthos on oxygen and nutrient fluxes between eulittoral sediments and associated water phases in the Elbe estuary. In: Colombo G, Ferrari I, Ceccherelli VU, and Rossi R (eds.) Marine Eutrophication and Population Dynamics, pp. 63--70. Fredensborg: Olsen & Olsen. Zevenboom W (1993) Assessment of eutrophication and its effects in marine waters. German Journal of Hydrography Suppl 1: 141--170.


EVAPORATION AND HUMIDITY K. Katsaros, Atlantic Oceanographic and Meteorological Laboratory, NOAA, Miami, FL, USA Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 870–877, & 2001, Elsevier Ltd.

Introduction Evaporation from the sea and humidity in the air above the surface are two important and related aspects of the phenomena of air–sea interaction. In fact, most subsections of the subject of air–sea interaction are related to evaporation. The processes that control the flux of water vapor from sea to air are similar to those for momentum and sensible heat; in many contexts, the energy transfer associated with evaporation, the latent heat flux, is of greatest interest. The latter is simply the internal energy carried from the sea to the air during evaporation by water molecules. The profile of water vapor content is logarithmic in the outer layer, from a few centimeters to approximately 30 m above the sea, as it is for wind speed and air temperature under neutrally stratified conditions. The molecular transfer rate of water vapor in air is slow and controls the flux only in the lowest millimeter. Turbulent eddies dominate the vertical exchange beyond this laminar layer. Modifications to the efficiency of the turbulent transfer occur due to positive and negative buoyancy forces. The relative importance of mechanical shear-generated turbulence and density-driven (buoyancy) fluxes was formulated in the 1940s, the Monin-Obukhov theory, and the field developed rapidly into the 1960s. New technologies, such as the sonic anemometer and Lyman-alpha hygrometer, were developed, which allowed direct measurements of turbulent fluxes. Furthermore, several collaborative international field experiments were undertaken. A famous one is the ‘Kansas’ experiment, whose data were used to formulate modern versions of the ‘flux profile’ relations, i.e., the relationship between the profile in the atmosphere of a variable such as humidity, and the associated turbulent flux of water vapor and its dependence on atmospheric stratification. The density of air depends both on its temperature and on the concentration of water vapor. Recent improvements in measurement techniques and the ability to measure and correct for the motion of a ship or aircraft in three dimensions have allowed more direct measurements of evaporation over the ocean. The fundamentals of turbulent transfer in the

324

atmosphere will not be discussed here, only the special situations that are of interest for evaporation and humidity. As the water molecules leave the sea, they remove heat and leave behind an increase in the concentration of sea salts. Evaporation, therefore, changes the density of salt water, which has consequences for water mass formation and general oceanic circulation. This article will focus on how humidity varies in the atmosphere, on the processes of evaporation, and how it is modified by the other phenomena discussed under the heading of air–sea interaction. All processes occurring at the air–sea interface interact and modify each other, so that none are simple and linear and most result in feedback on the phenomenon itself. The role of wind, temperature, humidity, wave breaking, spray, and bubbles will be broached and some fundamental concepts and equations presented. Methods of direct measurements and estimation using in situ mean measurements and satellite measurements will be discussed. Subjects requiring further research are also explored.

History/Definitions and Nomenclature Many ways of measuring and defining the quantity of the invisible gas, water vapor, in the air have developed over the years. The common ones have been gathered together in Table 1, which gives their name, definition, SI units, and some further explanations. These quantitative definitions are all convertable one into another. The web-bulb temperature may seem rather anachronistic and is completely dependent on a rather crude measurement technique, but it is still a fundamental and dependable measure of the quantity of water vapor present in the air. Evaporation or turbulent transfer of water vapor in the air was first modeled in analogy with down– gradient transfer by molecular conduction in solids. The conductivity was replaced by an ‘Austaush’ coefficient, Ae, or eddy diffusion coefficient, leading to the expression: E ¼ Ae r

@ q¯ @z

½1

where E is the evaporation rate, r the air density, q¯ is mean atmospheric humidity, and z represents the vertical coordinate. Assuming no advection, steady state, and no accumulation of water vapor in the surface layer of the atmosphere (referred to as ‘the


EVAPORATION AND HUMIDITY

Table 1

325

Measures of humidity

Nomenclature

Units (SI)

Definition

Absolute humidity Specific humidity

kg m3 g kg1

Mixing ratio

g kg1

Saturation humidity

Any of the above units

Relative humidity (RH) Vapor pressure Dew point temperature

% hPa (or mb) 1K, 1C

Wet-bulb temperature

1K, 1C

Amount of water vapor in the volume of associated moist air The mass of water per unit mass of moist air (or equivalently in the same volume) The ratio of the mass of water as vapor to the mass of dry air in the same volume Can be given in terms of all three units and refers to the maximum amount the air can hold at its current temperature in terms of absolute or specific humidity, corresponds to 100% relative humidity Percent of saturation humidity that is actually in the air The partial pressure of the water vapor in the air The temperature at which dew would form based on the actual amount of water vapor in the air. Dew point depression compared to actual temperature is a measure of the ‘dryness’ of the air This is a temperature obtained by the wetted thermometer of the pair of thermometers used in a psychrometera (see Measurements chapter)

a A psychrometer is a measuring device consisting of two thermometers (mercury in glass or electronic), where one thermometer is covered with a wick wetted with distilled water. The device is aspirated with environmental air (at an air speed of at least 3 m s1). The evaporation of the distilled water cools the air passing over the wet wick, causing a lowering of the wet thermometer’s temperature, which is dependent on the humidity in the air.

constant flux layer’), the Ae is a function of z as the turbulence scales increase away from the air–sea interface and the gradient is a decreasing function of height, z, as the distance from the source of water vapor, the sea surface, increases. Determining E by measuring the gradient of q has not proved to be a good method because of the difficulties of obtaining differences of q accurately enough and in knowing the exact heights of the measurements well enough (say from a ship or a buoy on the ocean). The Ae must also be determined, which would require measurements of the intensity of the turbulent exchange in some fashion. The socalled direct method for evaluating the vapor flux in the atmosphere requires high frequency measurements. This method has been refined during the past 35 years or so, and has produced very good results for the turbulent flux of momentum (the wind stress). Fewer projects have been successful in measuring vapor flux over the ocean, because the humidity sensors are easily corrupted by the presence of spray or miniscule salt particles on the devices, which being hygroscopic, modify the local humidity. Evaporation, E, can be measured directly today by obtaining the integration over all scales of the turbulent flux, namely, the correlation between the deviations from the mean of vertical velocity (w0 ) and humidity (q0 ) at height (z) within the constant flux layer. This correlation, resulting from the averaging of the vapor conservation equation (in analogy to the Reynolds stress term in the Navier–Stokes equation) can be measured directly, if sensors are available that resolve all relevant scales of fluctuations.

The correlation equation is rw q ¼ r¯ w ¯ q¯ þ rw ¯ 0 q0 ;

½2

where w and q are the instantaneous values and the overbar indicates the time-averaged means. The product of the averages is zero since w ¼ 0. Much discussion and experimentation has gone into determining the time required to obtain a stable mean value of the eddy flux rq ¯ 0 w0 . For the correlation term 0 0 rw ¯ q to represent the total vertical flux, there has to be a spectral gap between high and low frequencies of fluctuations, and the assumption of steady state and horizontal homogeneity must hold. The required averaging time is of the order of 20 min to 1 h. Another commonly used method, the indirect or inertial dissipation method, also requires high frequency sensing devices, but relies on the balance between production and destruction of turbulence to be in steady state. The dissipation is related to the spectral amplitude of turbulent fluctuations in the inertial subrange, where the fluctuations are broken down from large-scale eddies to smaller and smaller scales, which happens in a similar fashion regardless of scale of the eddies responsible for the production of turbulence in the atmospheric boundary layer. The magnitude of the spectrum in the inertial subrange is, therefore, a measure of the total energy of the turbulence and can be interpreted in terms of the turbulent flux of water vapor. The advantage of this method over the eddy correlation method is that it is less dependent on the corrections for flow distortion and motion of the ship or the buoy platform, but it


326


requires corrections for atmospheric stratification and other predetermined coefficients. It would not give the true flux if the production of turbulence was changing, as it does in changing sea states. Most of the time, the direct flux is not measured by either the direct or the indirect method; we resort to a parameterization of the flux in terms of so-called ‘bulk’ quantities. The bulk formula has been found from field experiments where the total evaporation E has been measured directly together with mean values of q and wind speed, U, at one height, z ¼ a (usually referred to as 10 m by adjusting for the logarithmic vertical gradient), and the known sea surface temperature. E ¼ rw0 q0 ¼ r¯ CEa Ua ðqs qa Þ

½3

where qs is the saturation specific humidity at the air– sea interface, a function of sea surface temperature (SST). Air in contact with a water surface is assumed to be saturated. Above sea water the saturated air has 98% of the value of water vapor density at saturation over a freshwater surface, due to the effects of the dissolved salts in the sea. CEa is the exchange

coefficient for water vapor evaluated for the height a. Experiments have shown CEa to be almost constant at 1.1–1.2 103 for Uo18 m s1, for neutral stratification, i.e. no positive or negative buoyancy forces acting and at a height of 10 m, written as CE10N. However, measurements show large variability in CE10N which may be due to the effects of sea state, such as sheltering in the wave troughs for large waves and increased evaporation due to spray droplets formed in highly forced seas with breaking waves. Results from a field experiment, the Humidity Exchange Over the Sea (HEXOS) experiment in the North Sea, are shown in Figure 1. Its purpose was to address the question of what happens to evaporation or water (vapor) flux at high wind speeds. However, the wind only reached 18 m s1 and the measurements showed only weak, if any, effects of the spray. Theories suggest that the effects will be stronger above 25 m s1. More direct measurements are still required before these issues can be settled, especially for wind speeds 420 m s1 (see Further Reading and the section on meteorological sensors for mean measurements for a discussion of the difficulties of making measurements over the sea at high wind speeds).

3

3

10 CEN

2

1

0 0

5

10

15

20

25

_

U10N (m s 1) Figure 1 Vapor flux exchange coefficients from two simultaneous measurement sets: the University of Washington (crosses) and Bedford Institute of Oceanography (squares) data. Thick dashed line is the average value, 1.12 103, for 170 data points. Thin dashed lines indicate standard deviations (from DeCosmo et al., 1996).



Clausius–Clapeyron Equation The Clausius–Clapeyron equation relates the latent heat of evaporation to the work required to expand a unit mass of liquid water into a unit mass of water as vapor. The latent heat of evaporation is a function of absolute temperature. The Clausius–Clapeyron equation expresses the dependence of atmospheric saturation vapor pressure on temperature. It is a fundamental concept for understanding the role of evaporation in air–sea interaction on the large scale, as well as for gaining insight into the process of evaporation from the sea (or Earth’s) surface on the small scale. Note first of all that the Clausius–Clapeyron equation is highly non-linear, viz: d ln rv DHvap ¼ dT RT 2

½4

where pv is the vapor pressure, T is absolute temperature (1K), and DHvap is the value of the latent heat of evaporation, R is the gas constant for water vapor ¼ 461.53 J kg1 1K1. The dependence of vapor pressure on temperature is presented in a simplified form as: T0 ðP a Þ es ¼ 610:8 exp 19:85 1 T

½5

where es is vapor pressure in pascals, T0 is a reference temperature set to 01C ¼ 273.16 1K, and T is the actual

327

temperature in 1K which is accurate to 2% below 301C (Figure 2). Figure 2 displays the saturation vapor pressure and the pressure of atmospheric water vapor for 60% relative humidity. On the right-hand side of the figure, the ordinate gives the equivalent specific humidity values (for a near surface total atmospheric pressure of 1000 hpa). This figure illustrates that the atmosphere can hold vastly larger amounts of water as vapor at temperatures above 201C than at temperatures below 101C. For constant relative humidity, say 60%, the difference in specific humidity or vapor pressure in the air compared with the amount at the air–sea interface, if the sea is at the same temperature as the air, is about three times at 301C what it would be at 101C. Therefore, evaporation is driven much more strongly at tropical latitudes compared with high latitudes (cold sea and air) for the same mean wind and relative humidity as illustrated by eqn [4] and Figure 2.

Tropical Conditions of Humidity By far, most of the water leaving the Earth’s surface evaporates from the tropical oceans and jungles, providing the accompanying latent heat as the fuel that drives the atmospheric ‘heat engines,’ namely, thunderstorms and tropical cyclones. Such extreme and violent storms depend for their generation on the enormous release of latent heat in clouds to create the vertical motion and compensating horizontal

Vapor pressure vs temperature 100

Vapor pressure (hPa)

80

60

40

20

0 0

5

10

15

20

25

30

40

Temperature (˚C) 60% RH

Saturation /100% RH

Figure 2 Vapor pressure (hPa) as a function of temperature for two values of relative humidity, 60% and 100%.


45

328


accelerated inflows. Tropical cyclones do not form over oceanic regions with temperatures o261C, and temperature increases of only 11 or 21C sharply enhance the possibility of formation.

Latitudinal and Regional Variations The Clausius–Clapeyron equation holds the secrets to the role of water vapor for both weather and climate. Warm moist air flowing north holds large quantities of water. As the air cools by vertical motion, contact with cold currents, and loss of heat by infrared radiation, the air reaches saturation and either clouds, storms and rain form, or fog (over cold surfaces) and stratus clouds. The warmer and moister the original air, the larger the possible rainfall and the larger the release of latent heat. Latitudinal, regional, and seasonal variations in evaporation and atmospheric humidity are all related to the source of heat for evaporation (upper ocean heat content) and the capacity of the air to hold water at its actual temperature. Many other processes such as the dynamics behind convergence patterns and the development of atmospheric frontal zones contribute to the variability of the associated weather.

Vertical Structure of Humidity The fact that the source of moisture is the ocean, lakes, and moist ground explains the vertical structure of the moisture field. Lenses of moist air can form aloft. However, when clouds evaporate at high elevations where atmospheric temperature is low, the absolute amounts of water vapor are also low for that reason. Thus, when the surface air is continually mixed in the atmospheric boundary layer with drier air, being entrained from the free atmosphere across the boundary layer inversion, it usually has a relative humidity less than 100% of what it could hold at its actual temperature. The exceptions are fog, clouds, or heavy rain, where the air has close to 100% relative humidity. The process of exchange between the moist boundary layer air and the upper atmosphere allows evaporation to continue. Deep convection in the inter-tropical convergence zone brings moist air up throughout the whole of the troposphere, even over-shooting into the stratosphere. Moisture that does not rain out locally is available for transport poleward. The heat released in these clouds modifies the temperature of the air. Similarly, over the warm western boundary currents, such as the Gulf Stream, Kuroshio, and Arghulas Currents, substantial evaporation and warming of the

atmosphere takes place. Without the modifying effects of the hydrologic cycle of evaporation and precipitation on the atmosphere, the continents would have more extreme climates and be less habitable.

Sublimation–Deposition The processes of water molecules leaving solid ice and condensing on it are called sublimation and deposition, respectively. These processes occur over the ice-covered polar regions of the ocean. In the cold regions, this flux is much less than that from open leads in the sea ice due to the warm liquid water, even at 01C. At an ice surface, water vapor saturation is less than over a water surface at the same temperature. This simple fact has consequences for the hydrologic cycle, because in a cloud consisting of a mixture of ice and liquid water particles, the vapor condenses on the ice crystals and the droplets evaporate. This process is important in the initial growth of ice particles in clouds until they become large enough to fall and grow by coalescence of droplets or other ice crystals encountered in their fall. Similar differences in water vapor occur for salty drops, and the vapor pressure over a droplet also depends on the curvature (radius) of the drop. Thus, particle size distribution in clouds and in spray over the ocean are always changing due to exchange of water vapor. For drops to become large enough to rain out, a coalescencetype growth process must typically be at work, since growth by condensation is rather slow.

Sources of Data Very few direct measurements of the flux of water vapor are available over the ocean at any one time. The mean quantities (U, qa , SST) needed to evaluate the bulk formula are reported regularly from voluntary observing ships (VOS) and from a few moored buoys. However, most of such buoys do not measure surface humidity, only a small number in the North Atlantic and tropical Pacific Oceans do so. The VOS observations are confined to shipping lanes, which leaves a huge void in the information available from the Southern Hemisphere. Alternative estimates of surface humidity and the water vapor flux include satellite methods and the surface fluxes produced in global numerical models, in particular, the re-analysis projects of the US Weather Service’s National Center for Environmental Prediction (NCEP) and the European Center for Medium Range Weather Forecasts (ECMWF). The satellite method has large statistical uncertainty and,



thus, requires weekly to monthly averages for obtaining reasonable accuracy (730 W m2 and 715 W m2 for the weekly and monthly latent heat flux). Therefore, these data are most useful for climatological estimates and for checking the numerical models’ results.

Estimation of Evaporation by Satellite Data The estimation of evaporation/latent heat flux from the ocean using satellite data also relies on the bulk formula. The computation of latent heat flux by the bulk aerodynamic method requires SST, wind speed (U10N ), and humidity at a level within the surface layer qa , as seen in eqn [3]. Therefore, evaluation of the three variables from space is required. Over the ocean, U10N and SST have been directly retrieved from satellite data, but qa has not. A method of estimating qa and latent heat flux from the ocean using microwave radiometer data from satellites was proposed in the 1980s. It is based on an empirical relation between the integrated water vapor W (measured by spaceborne microwave radiometers) and qa on a monthly timescale. The physical rationale is that the vertical distribution of water vapor through the whole depth of the atmosphere is coherent for periods longer than a week. The relation does not work well at synoptic and shorter timescales and also fails in some regions during summer. Modification of this method by including additional geophysical parameters has been proposed with some overall improvement, but the inherent limitation is the lack of information about the vertical distribution of q near the surface. Two possible improvements in E retrieval include obtaining information on the vertical structures of humidity distribution and deriving a direct relation between E and the brightness temperatures (TB) measured by a radiometer. Recent developments provide an algorithm for direct retrieval of boundary layer water vapor from radiances observed by the Special Sensor Microwave/Imager (SSM/I) on operational satellites in the Defense Meteorological Satellite Program since 1987. This sensor has four frequencies, 19.35, 22, 37, and 85.5 GHz, all except the 22 GHz operated at both horizontal and vertical polarizations. The 22 GHz channel at vertical polarization is in the center of a weak water vapor absorption line without saturation, even at high atmospheric humidity. The measurements are only possible over the oceans, because the oceans act as a relatively uniform reflecting background. Over land, the signals from the ground overwhelm the water vapor information.

329

Because all the three geophysical parameters, U10N , W, and SST, can be retrieved from the radiances at the frequencies measured by the older microwave radiometer, launched in 1978 and operating to 1985 – the Scanning Multichannel Microwave Radiometer (SMMR) on Nimbus-7 (similar to SSM/I, but with 10.6 and 6.6 GHz channels as well, and no 85 GHz channels) – the feasibility of retrieving E directly from the measured radiances was also demonstrated. SMMR measures at 10 channels, but only six channels were identified as significantly useful in estimating E. SSM/I, the operational microwave radiometer that followed SMMR, lacks the low-frequency channels which are sensitive to SST, making direct retrieval of E from TB unfeasible. The microwave imager (TMI) on the Tropical Rainfall Measuring Mission (TRMM), launched in 1998, includes low-frequency measurements sensitive to SST and could, therefore, allow direct estimates of evaporation rates. Figure 3 gives an example of global monthly mean values of humidity obtained solely with satellite data from SSM/I. To calculate qs, gridded data of sea surface temperature can also be used, such as those provided operationally by the US National Weather Service based on infrared observations from the Advanced Very High Resolution Radiometer (AVHRR) on operational polar-orbiting satellites. The exact coincident timing is not so important for SST, since SST varies slowly due to the large heat capacity of water, and this method can only provide useful accuracies when averages are taken over 5 days to a week. Wind speed is best obtained from scatterometers, rather than from the microwave radiometer, in regions of heavy cloud or rain, since scatterometers (which are active radars) penetrate clouds more effectively. Scatterometers have been launched in recent times by the European Space Agency (ESA) and the US National Aeronautic and Space Administration (NASA) (the European Remote Sensing Satellites 1 and 2 in 1991 and 1995, the NASA scatterometer, NSCAT, on a Japanese short-lived satellite in 1996, and the QuikSCAT satellite in 1999).

Future Directions and Conclusions Evaporation has been measured only up to wind speeds of 18 m s1. The models appear to converge on the importance of the role of sea spray in evaporation, indicating that its significance grows beyond about 20 m s1. However, the source function of spray droplets as a function of wind speed or wave breaking has not been measured, nor are techniques for measuring evaporation in the


330


_ 50

0

50

100

150

200

250

Figure 3 Global distribution of monthly mean latent heat flux in W m2 for September 1987. (Reproduced with permission from Schulz et al., 1997.)

presence of droplets well–developed, whether for rain or sea spray. Since evaporation and the latent heat play such important roles in tropical cyclones and many other weather phenomena, as well as in oceanic circulation, there is great motivation for getting this important energy and mass flux term right. The bulk model is likely to be the main method used for estimating evaporation for some time to come. Development of more direct satellite methods and validating them should be an objective for climatological purposes. Progress in the past 30 years has brought the estimate of evaporation on a global scale to useful accuracy.

See also IR Radiometers. Satellite Remote Sensing of Sea Surface Temperatures. Sensors for Mean Meteorology.

Further Reading Bentamy A, Queffeulou P, Quilfen Y, and Katsaros KB (1999) Ocean surface wind fields estimated from satellite active and passive microwave instruments. Institute of Electrical and Electronic Engineers, Transactions, Geoscience Remote Sensing 37: 2469--2486. Businger JA, Wyngaard JC, lzumi Y, and Bradley EF (1971) Flux-profile relationships in the atmospheric surface layer. Journal of Atmospheric Science 28: 181--189.

DeCosmo J, Katsaros KB, Smith SD, et al. (1996) Air–sea exchange of water vapor and sensible heat: The Humidity Exchange Over the Sea (HEXOS) results. Journal of Geophysical Research 101: 12001--12016. Dobson F, Hasse L, and Davies R (eds.) (1980) Instruments and Methods in Air–sea Interaction. New York: Plenum Publishing. Donelan MA (1990) Air–sea Interaction. In: LeMehaute B and Hanes DM (eds.) The Sea, Vol. 9, pp. 239--292. New York: John Wiley. Esbensen SK, Chelton DB, Vickers D, and Sun J (1993) An analysis of errors in Special Sensor Microwave Imager evaporation estimates over the global oceans. Journal of Geophysical Research 98: 7081--7101. Geernaert GL (ed.) (1999) Air–sea Exchange Physics, Chemistry and Dynamics. Dordrecht: Kluwer Academic Publishers. Geernaert GL and Plant WJ (eds.) (1990) Surface Waves and Fluxes, Vol. 2. Dordrecht: Kluwer Academic Publishers. Katsaros KB, Smith SD, and Oost WA (1987) HEXOS – Humidity Exchange Over the Sea: A program for research on water vapor and droplet fluxes from sea to air at moderate to high wind speeds. Bulletin of the American Meteoroloical Society 68: 466--476. Kraus EB and Businger JA (eds.) (1994) Atmosphere– Ocean Interaction 2nd ed. New York: Oxford University Press. Liu WT and Katsaros KB (2001) Air–sea fluxes from satellite data. In: Siedler G, Church J, and Gould J (eds.) Ocean Circulation and Climate. Academic Press Liu WT, Tang W, and Wentz FJ (1992) Precipitable water and surface humidity over global oceans from SSM/I and ECMWF. Journal of Geophysical Research 97: 2251--2264.



Makin VK (1998) Air–sea exchange of heat in the presence of wind waves and spray. Journal of Geophysical Research 103: 1137--1152. Schneider SH (ed.) (1996) Encyclopedia of Climate and Weather. New York: Oxford University Press. Schulz J, Meywerk J, Ewald S, and Schlu¨ssel P (1997) Evaluation of satellite-derived latent heat fluxes. Journal of Climate 10: 2782--2795.

331

Smith SD (1988) Coefficients for sea surface wind stress, heat flux, and wind profiles as a function of wind speed and temperature. Journal of Geophysical Research 93: 15467--15472.


EXOTIC SPECIES, INTRODUCTION OF D. Minchin, Marine Organism Investigations, Killaloe, Republic of Ireland

History

& 2009 Elsevier Ltd. All rights reserved.

Species have been moved for several hundreds of years and some have almost certainly been carried with the earliest human expeditions. Evidence for early movement is scant; this can normally be determined only from hard remains, from a wellunderstood biogeography of a taxonomic group or perhaps from genetic comparisons. Although the soft-shell clam Mya arenaria was present in northern Europe during the late Pliocene, it disappeared during the last glacial period. In the thirteenth century their shells were found in north European middens; this was considered evidence of its introduction by returning Vikings from North America. It may have been used as fresh food during long sea journeys, as it is easily collected and perhaps was maintained in the bilges of their vessels, or may have unintentionally been carried with sediment used as ballast discharged on return to Europe. Certainly it became sufficiently abundant in the Baltic at about this time to be referred to by quaternary geologists as the Mya Sea. This species was also introduced to the Black Sea in the 1960s and rapidly became abundant, resembling its sudden apparent Baltic expansion. The South American coral Oculina patagonica, established in southern Europe, was most probably introduced on sailing ship hulls returning from South America during the sixteenth or seventeenth centuries. Indeed these sailing vessels probably carried a wide range of organisms attached to their hulls. Predictions of the most likely fouling species during these times are based on settlements on panels attached to the hull of a reconstructed sailing ship. Hulls of wooden vessels were fouled with complex communities and excavated by boring organisms whose vacant galleries could provide refugia for a wide range of species including mobile animals. The drag imposed by fouling and the structural damage caused by the boring organisms played its role in the outcome of naval engagements and in the duration of the working life of these vessels, particularly in tropical regions where boring activities and fouling communities evolve rapidly. The periods of stay by sailing ships in ports could provide opportunities for creating new populations arising from drop-off or from reproduction while remaining attached. Some vessels never returned but would have provided instant reefs by sinking, wreckage, or abandonment, thereby distributing a wide assemblage of exotic species. Wooden vessels are still in widespread use

Introduction Exotic species, often referred to as alien, nonnative, nonindigenous, or introduced species, are those that occur in areas outside of their natural geographic range. Vagrant species are those that appear from time to time beyond their normal range and are often confused with exotic species. Since marine science evolved following periods of human exploration and worldwide trade, there are species that may have become introduced at an early time but cannot clearly be ascribed as native or exotic; these are known as cryptogenic species. The full contribution of exotic species among native assemblages remains, and probably will continue to remain, unknown, but these add to the diversity of an area. There are no documented accounts of an introduced species resulting in the extinction of native species in marine habitats as has occurred in freshwater systems. Nevertheless, exotic species can result in habitat modifications that may reduce native species abundance and restructure communities. The greatest numbers of exotic species are inadvertently distributed by shipping either attached to the hull or carried in the large volumes of ballast water used for ship stability. Introductions may also be deliberate. The dependence for food in developing countries and expansion of luxury food products in the developed world have led to increases in food production by cultivation of aquatic plants, invertebrates, and fishes. Many native species do not perform as well as the desired features of some introduced organisms now in widespread cultivation, for example, the Pacific oyster Crassostrea gigas and Atlantic salmon Salmo salar. Unfortunately, unwanted organisms that have been unintentionally introduced with stock movements can reduce production. Consequently, care is needed when introductions are made. An accidentally introduced exotic species may remain unnoticed until such time as it either becomes abundant or causes harmful effects, whereas larger organisms are normally recognized sooner. Little is known about the movement of the smallest organisms, yet these must be in transit in great numbers every day.

332


333

EXOTIC SPECIES, INTRODUCTION OF

and continue to endure problems of drag and structural damage. Many preparations have been used for controlling this, with the most effective being developed over the last 150 years. Wooden vessels became sheathed with thin copper plates and were vulnerable only where these became displaced or damaged. Ironclad sailing vessels also evolved and normal hull life became prolonged because damage was considerably reduced. The building of iron and then steel ship hulls eliminated opportunities for boring and cryptic fauna, yet hull surfaces needed protection from rusting and fouling. The development of protective and also toxic coatings then evolved. The incorporation of copper and other salts in these coatings considerably reduced fouling, and therefore drag, and protected the underlying surfaces from corrosion. Antifouling coatings, in particular, organotins, such as tributyltin compounds, have been very effective in fouling control, and have enabled most vessels to remain in operation before reentering dry dock for servicing, about every 5 years. Unfortunately, tributyltin is a powerful biocide harmful to aquatic organisms in port regions. Its use during the 1980s and 1990s became restricted or banned in many countries on vessels o25 m. And between 2003 and 2008 all merchant shipping are expected to phase out its use. New products are being developed with an emphasis on less toxic and nontoxic applications. Most exotic species, used in mariculture or for developing fisheries, were spread following the development of steam transport. An early example is the transport from the eastern coast of North America of the striped bass Morone saxatilis fingerlings to California in the 1880s by train. Within 10 years, it became regularly sold in the fish markets of San Francisco. Attempts at culturing different species developed slowly due to a poor knowledge of the full life cycle. Planktonic stages were especially misunderstood and pond systems for rearing oyster larvae, for example, were not fully successful until the larval biology became fully known. There were strong economic pressures to develop this knowledge at an early stage. The progressive understanding of behavior and the physical requirements of organisms, and development of algae, brine shrimp, and other cultures and the use of synthetic materials, such as plastics, enabled a rapid expansion of marine culture products ranging from pearls, chemical compounds, to food.

The Vectors of Exotic Species Organisms are deliberately introduced for culture, to create fisheries, or as ornamental species. Future

2

4

1 3

Figure 1 Some examples of primary inoculations: 1, Carcinus maenas, the shore crab, a predator of mollusks (vector, probably with shore macroalgae used for transporting fishing baits from the North Atlantic); 2, Mnemiopsis leidyi a comb jelly, planktonic predator of larval fishes and crustaceans (vector, ballast water); 3, Styela clava an Asian tunicate that causes trophic competition by filtering the water in docks and estuaries (vector, hull fouling); 4, Asterias amurensis, an Asian sea star, avid predator of mollusks (vector, ballast water or hull fouling).

introductions may involve species producing products used in pharmacy, as food additives, in the management of diseases, for biological control, or for water quality management. A primary inoculation is where the first population of an exotic becomes established in a new biogeographic region (Figure 1). A secondary inoculation arises by means of further established populations from the first population to other nearby regions by a wide range of activities and so the risk of spread is increased. In many cases, these range extensions are inevitable, either because of human behavior and marketing patterns or because of their mode of life. For example, the Indo-Pacific tubeworm Ficopomatus enigmaticus can extensively foul brackish areas and once transported by shipping is easily carried to other estuaries by leisure craft or aquaculture activities, or both. Species are introduced by a wide range of vectors that can include attachment to seaweeds used for packing material or releases of imported angling baits. The principal vectors are discussed below. Aquaculture

Useful species have a high tolerance of handling, can endure a wide range of physical conditions, are easily induced to reproduce, and have high survival throughout their production phase. Because of capital costs to retain stock in cages and/or trays, the species must, most usually, be cultivated at densities higher than occurs naturally. The species most likely to adapt to these conditions will be those living under variable conditions at high densities, such as mussels and oysters and some shoaling fishes. Species


334


that do not naturally thrive under these conditions may require more attention and will tend to depend on speciality markets, such as some scallops, tuna, and grouper. Further, aquaculture efforts target those species which grow fast to marketable size. Mollusks High-value species that are tolerant to handling as well as a wide range of physical conditions are preferred. Some have a long shelf life that allows for live sales. Mollusks, such as clams (e.g., Ruditapes philippinarum, Mercenaria mercenaria), scallops (e.g., Patinopecten yessoensis, Argopecten irradians), oysters (e.g., Ostrea edulis, C. gigas), and abalone (e.g., Haliotis discus hannai, Haliotis rufescens), have been introduced for culture (Figure 2). Because oysters survive under cool damp conditions for several days, large consignments are easily transported long distances and so have become widely distributed since the advent of steam transportation. With the exception of the east coast of North America, the Pacific oyster C. gigas, for example, is now widespread in the Northern Hemisphere and accounts for 80% of world oyster production. Its wide tolerance of salinity, temperature, and turbid conditions and rapid growth make it a desirable candidate for culture. In many cases, it has replaced the native oyster production where this declined either because of overexploitation, diseases of former stocks, or to develop new culture areas. However, pests, parasites, and diseases can be carried within the tissues, mantle cavity, and both in and on the shells of mollusks when they are moved, and can compromise molluskan culture and wild fisheries and may also modify ecosystems. Some examples of pests moved with oysters include the slipper limpet Crepidula fornicata and the sea squirt Styela clava (Figure 2). Crustaceans The worldwide expansion of penaeid shrimp culture has led to a series of unregulated movements resulting in serious declines of shrimp production caused by pathogenic viruses, bacteria, protozoa, and fungi. Viruses have caused the most serious mortalities of farmed shrimps’ broodstock and pose a deterrent to developing culture projects because of casual broodstock movements rather than using those certified as specific-pathogen-free. One parvovirus is the infectious hypodermal and hematopoietic necrosis virus found in wild juvenile and adult prawn stages. This virus was endemic to Southeast Asia, Indonesia, and the Philippines and is now widespread at farms in the tropical regions of the Indo-Pacific and the Americas. About six serious viruses of penaeid shrimp are known. Production may be limited unless disease-resistant

stocks for viruses can be developed together with vaccination against bacterial diseases. Fishes Of the thousands of fish species only a small number generate a market price high enough to cover production costs. Such fish need to have a high flesh to body weight and high acceptance. While in culture, they should be tolerant of handling and to the wide range of seasonal and diurnal conditions and should not be competitive. The North Atlantic salmon S. salar is one of the very few exotics in culture in both hemispheres in the cool to warm temperate climates of North America, Japan, Chile, New Zealand, and Tasmania (Figure 2). Fertilized eggs of improved stock are in constant movement between these countries. However, various diseases such as infectious salmon anemia (ISA) may limit the extent of future movements. It is likely that intensive shore culture facilities for marine species will become more common as the physiological and behavioral requirements of promising species become better understood. For example, in Japan, the bastard sole Paralichthys olivaceus is extensively cultivated in shore tanks (and has been introduced to Hawaii) and the sole Solea senegalensis in managed lagoons in Portugal. Cultivation under these circumstances is likely to lead to better control of pests, parasites, and diseases, whereas cage culture under more exposed conditions off shore may lead to unexpected mortalities from siphonophores, medusae, algal blooms, and epizootic infestations from parasites, some perhaps introduced. Shipping

Much of the world trade depends on shipping; the scale and magnitude of these vessels are seldom appreciated. To travel safely, they must either carry cargo or water (as ballast) so that the vessel is correctly immersed to provide more responsive steerage, by allowing better propulsion (without cavitation), rudder bite, and greater stability. The amount of ballast water carried can amount to 30% or more of the overall weight of the ship. Ships also have a large immersed surface area to which organisms attach and result in increased drag that results in higher fuel costs. Nevertheless, despite the best efforts of management, ships carry a great diversity and large numbers of species throughout the world. Unfortunately the risk of introducing further species increases because more ships are in transit and many travel faster than before and operate a wide trading network with new evolving commercial links.


336


Ships’ ballast water Ballast water is held on board in specially constructed tanks used only for holding water. The design and size of these tanks depend on the type and size of vessel (Figure 3). In some vessels that carry bulk products, a cargo hold may also be flooded with ballast water, but these vessels will also have ballast tanks. Ballast tanks are designed to add structural support to the ship, and will have access ladders, frames, and perforated platforms and baffles to reduce water slopping, making them complex engineering structures. All vessels carry ballast for trim; some of this may be permanent ballast (in which case it is not exchanged at any time) or it may be ‘all’ released (there is always a small portion of unpumpable water that remains) before taking on cargo. There are several tanks on a ship, and because there may be partial discharges of water, the water within nonpermanent ballast tanks can contain a mixture of water from different ports. Pumping large volumes of water when ballasting is time-consuming and costly. Should a vessel be

Bulk carrier

General cargo

Ore carrier

Tanker

Container

Roll-on roll-off

Ore carrier

Oil tanker

Figure 3 Ballast tanks in ships are dedicated for carrying water for the necessary stabilization of ships. The ballast tanks are outlined in black and show the variation of ballast tank position and size in relation to the type of vessel. Wing tanks (ballast tanks situated along the sides) are not shown in this diagram.

ballasted incorrectly, the structural forces produced could compromise the hull. Ships have broken up in port because of incorrect deballasting procedures in relation to the loading/unloading of cargo. Other vessels are known to have capsized due to incorrect ballast water operations. For these reasons, there are special guidelines for the correct ballasting of tanks in relation to cargo load and weather conditions. Ballasting of water normally occurs close to a port, often in an estuary or shallow harbor; this can include turbid water, sewage discharges, and pollutants. The particulates settle inside the tanks to form sediment accumulations that may be 30 cm in depth. Worms, crustaceans, mollusks, protozoa, and the resting stages of dinoflagellates, as well as other species, can live in these muds. Resting stages may remain dormant for months or years before ‘hatching’. Ballast sediments and biota are thus an important component of ballast water uptake adding to the diversity of carried organisms released in new regions through larval releases or from sediment resuspension following journeys with strong winds. Fresh ballast water normally includes a wide range of different animal and plant groups. However, most of these expire over time, so that on long journeys fewer species survive, and those that do survive, do so in reduced numbers. Because ballast water is held in darkness the contained phytoplankton are unable to photosynthesize. Animals, dependent on these microscopic plants, during vulnerable stages in their development may expire because of insufficient food, unless the tanks are frequently flushed. Scavengers and predators may have better opportunities to survive. To date, observations of organisms surviving in ballast water are incomplete because of the inability or poor efficiency in obtaining adequate samples. Some ballast tanks may only be sampled through a narrow sounding pipe to provide a limited sample of the less active species from one region near the tank floor. Nevertheless, all studies to date point to a wide range of organisms surviving ballast transport, ranging from bacteria to adult fishes. The invasions of the comb jelly Mnemiopsis leidyi (from the eastern coast of North America to the Black Sea), the Asian sea star Asterias amurensis (from Japan to Tasmania), and the shore crab Carcinus maenas (to Australia, the Pacific and Atlantic coasts of North America, South Africa, and the Patagonian coast from Europe) probably arose from ballast water transport (Figure 2). Ships’ hulls Ships are dry-docked for inspection, structural repairs, and recoating the hull with antifouling paint about every 3–5 years. The interdocking time varies according to the type of



vessel and its age. While in dry dock, the ship is supported on wooden blocks. The areas beneath the blocks, which may in total cover a ship surface of >100 m2, are not painted and here fouling may freely develop once the ship is returned to service. The hull is cleaned by shot blasting or using powerful water jets, then recoated. Many coatings have a toxic surface that over time becomes leached or worn from the hull. Unfortunately, some toxic compounds, once released, have caused unexpected and unwanted environmental effects. The use of organotin paints such as tributyltin effectively reduced hull fouling since the 1970s, but it has caused sexual distortion in snails and reproductive and respiratory impairment in other taxa, especially in coastal waters along highly frequented shipping routes. Vessels trading in tropical waters have a higher rate of loss of the active coat and on these vessels a greater fouling can occur at an earlier time. Most organisms carried on ships’ hulls are not harmful but may act as the potential carriers of a wide range of pests, parasites, and diseases. Oysters and mussels frequently attach to hulls and may transmit their diseases to aquaculture sites in the vicinity of shipping ports. Small numbers of disease organisms, once released, may be sufficient to become established if a suitable host is found. Should the host be a cultivated species, it may rapidly spread to other areas in the course of normal trade before the disease becomes recognized. Many marine organisms spawn profusely in response to sea temperature changes. Spawning could arise following entry of a ship to a port, during unloading of cargo or ballast in stratified water, or from diurnal temperature changes. A ship may turn around in port in a few hours to leave behind larvae that may ultimately settle to form a new population. Vessels moored for long periods normally acquire a dense and complex fouling community that includes barnacles. Once of sufficient size, these can provide toxic-free surfaces for further attachment, in particular for other barnacles (e.g., the Australasian barnacle Elminius modestus, present in many Northern European ports), mollusks, hydroids, bryozoa, and tunicates (e.g., the Asian sea squirt S. clava, widely distributed in North America and northern Europe (Figure 2)). Vacant shells of barnacles and oysters cemented to the hull can provide shelter for mobile species such as crustaceans and nematodes, and some mobile species such as anemones and flatworms can attach directly to the hull. Small vessels such as yachts and motorboats also undertake long journeys and may become fouled and

337

thereby extend the range of attached biota. Normally small vessels would be involved in secondary inoculations from port regions to small inlets and lagoons (areas where shipping does not normally have access) and spread species such as the barnacle Balanus improvisus to other brackish areas. Aquarium Species

Aquarium releases seldom become established in the sea whereas in fresh water this is common. The attractive green feather-like alga Caulerpa taxifolia was probably released into the Mediterranean Sea from an aquarium in Monaco and by 2000 was present on the French Mediterranean coast, Italy, the Balearic Islands, and the Adriatic Sea. It has a ‘root’ system that grows over rock, gravel, or sand to form extensive meadows in shallows and may occur to depths of 80 m. It excludes seagrasses and most encrusting fauna and so changes community structure. Although it possesses mild toxins that deter some grazing invertebrates, several fishes will browse on it. This invasive plant is still expanding its range and may extend throughout much of the Mediterranean coastline and perhaps to some Atlantic coasts. Trade Agreements and Guaranteed Product Production

Trade agreements do not normally take account of the biogeographical regions among trading partners, so introductions of unwanted species are almost inevitable. Often, a population of a harmful species is further spread by trading and thereby has the ability to compromise the production of useful species. Trade in the same species from regions where quarantine was not undertaken may compromise cultured exotic species introduced through the expensive quarantine process which are disease-free. The risk of movement of serious diseases from one country to another is usually recognized, whereas pests are not normally regulated by veterinary regulations, and so are more likely to become freely distributed. Harvesting of cultured species may be prohibited in areas following the occurrence of toxin-producing phytoplankton. Toxins naturally occur in certain phytoplankton species (most usually dinoflagellates) and are filtered by the mollusk, and these become concentrated within molluskan tissues, with some organs storing greater amounts. Some of these toxins may subsequently accumulate within the tissues of molluskan-feeding crustaceans, such as crabs. If contaminated shellfish is consumed by humans, symptoms such as diahorritic shellfish poisoning, paralytic shellfish poisoning, neurological shellfish poisoning, and amnesic shellfish poisoning may



250 200

Port Phillip Bay San Francisco Bay Mediterrean Sea

150

Baltic Sea

100 50

1990

1980

1970

1960

1950

1940

1930

1920

1910

1900

1890

1880

0 1870

Natural eruptions of endemic or introduced pathogens may be responsible for chronic mortalities or unwanted phenomena. The die-off of the sea urchin Diadema antillarum in the western Atlantic may be due to a virus. The unsightly fibropapillomae of the green turtle Chelonia mydas, previously known only in the Atlantic Ocean, are now found in the IndoPacific, where it was previously unknown. The great mortalities of pilchard Sardinops sagax neopilchardus off Australia in the 1990s and mortalities of the bay scallop A. irradians in China may be a result of introduced microorganisms. In the case of introduced culture species, mortalities may ensue from a lack of resistance to local pathogens. Bonamia ostreae, a protozoan parasite in the blood of some oyster species, was unknown until found in the European flat oyster O. edulis following its importation to France from the American Pacific coast, where it had been introduced several years earlier. Very often harmful species first become noticed when aquaculture species are cultivated at high densities, for example, the rhizocephalan-like Pectinophilus inornata found in the scallop Pa. yessoensis in Japanese waters. Often careful studies of native biota reveal new species to science, such as the generally harmful protozoans, Perkensis species found in clams, oysters, and scallops that can be transmitted to each generation by adhering to eggs. Pests, parasites, and diseases in living products used in trade are likely sources of transmission. However, viruses or resting stages of some species may also be transmitted in frozen and dried products. Changes in global climate may be contributing to some of these appearances aided by high levels of human mobility.

Some areas favor exotic species with opportunities for establishment and so enable their subsequent spread to nearby regions. These areas are normally shipping ports within partly enclosed harbors with low tidal amplitudes and/or with good water retention and a large number of arriving vessels. Although it may be possible to predict which ports are the main sites for primary introductions, the factors involved are not clearly understood and information on the exotic species component is presently only available for a small number of ports. Nevertheless, ports with many exotic species are areas where further exotics will be found. Such regions are likely sites for introductions because ships carry a very large number of a wide range of species from different taxonomic groups. Port regions with known concentrations of exotic species include San Francisco Bay, Prince William Sound, Chesapeake Bay, Port Phillip Bay, Derwent Estuary, Brest Harbor, Cork Harbour, and The Solent (Figure 4). In the Baltic Sea, there are several ports that receive ballast water from ships operated via canals from the Black and Caspian Seas as well as arising from direct overseas trade. The component of the exotic species biomass in this region is high. In some areas such as the Curonian Lagoon, Lithuania, exotic species comprise the main biomass. The Black Sea has similar conditions to the Baltic Sea where established species have modified the economy of the region. The effects vary from dying clams Mya arenaria creating a stench on tourist beaches, poor recruitment of pilchard and anchovy due to a combination of high exploitation of the fisheries, changes in water quality and an increased predation of their larvae by an introduced comb jelly M. leidyi, to the development

1860

The Increasing Emergence of Unexplained Events

Vulnerable Regions

1850

occur. These conditions are caused by different toxins, and new toxins continue to be described. As they all have different breakdown rates in the shellfish tissues, there are varying periods when the products are prohibited from sale. There are monitoring programs for evaluating the levels of these contaminants in most countries. On occasions when harvesting of mollusks is suspended, consignments from abroad may be imported to maintain production levels. If these consignments arrive in poor condition, or are unsuitable because of heavy fouling, they may be relaid or dumped in the sea or on the shore. Such actions can extend the range of unwanted exotic species.

Number of exotics

338

Figure 4 Accumulative numbers of known exotic species in vulnerable areas. The real numbers of exotic species are probably much greater than shown.



of an industry based on harvesting and export of the large introduced predatory snail Rapana venosa. The Mediterranean Sea has increasing numbers of unintended introductions arising from trade, mainly by shipping, and from expansion of their range from the Red Sea via the Suez Canal (Figure 4). The majority of exotic species is found in temperate regions of the world. Whether this is a surmise because of a lack of a full understanding, or whether this is due to a real effect, is not known. It may be that tropical species are well dispersed because of natural vectors and that further transmission by shipping is of little consequence to the overall biota present. However, it may be that there is such a daunting diversity of species present in tropical regions, with many still to be described, that it is difficult to grasp the complexity and so be able to understand exotic species’ movements in this zone. In the contrasting colder climates, the likely slower development of species may inhibit an introduction from being successful. Temperate ports on either side of the same ocean appear to share several species in common, whereas species from temperate regions from other hemispheres or oceans are less common. This suggests that species that do not undergo undue physiological stresses and those present in shorter voyages have a greater probability of becoming established. For many species, although opportunities for their dispersal in the past may have taken place, their successful establishment has not succeeded. If an inoculation is to succeed, it must pass through a series of challenges. On most occasions, the populations are unable to be maintained at the point of release in sufficient numbers capable of establishment, even though some may survive. Any sightings of exotic individuals may often pose as a warning that this same species, under different conditions, could become established. The increasing volume and speed of shipping, and of air transport (in the case of foods for human consumption and aquarium species) will provide new opportunities. A successful transfer for a species will depend on: appropriate season, life-history stage in transit, survival time without food, re-immersion, temperature and salinity tolerance, rate of dispersal at the point of release, inoculation size and frequency, presence/absence of predators in the receiving environment, and many unknown factors.

Mode of Life The dispersal of a species from its point of introduction, and the speed at which it expands its range,

339

will depend on the behavior of its life-history stages and hydrodynamic conditions in the recipient habitat. Species with very short or no planktonic stages, and with limited mobility, are likely to remain close to the site of introduction, unless carried elsewhere by other means. Tunicates have short larval stages and a sessile adult life and so are normally confined to inlets. Buoyant and planktonic species, which include seaweeds with air bladders and many invertebrates, may be carried by combinations of wind and current and become rapidly dispersed in a directional pattern ruled by the principal vectors. Most species have dispersal potentials that lie between these extremes and some, such as active crustaceans and fishes, may become distributed over a wide range as a result of their own activities. The attempt to establish the pink salmon Onchorhynchus gorbuscha on the northern coast of Russia resulted in its capture as far south in the North Atlantic as Ireland. A recent and successful introduction of the king crab Paralithodes camtschaticus to the Barents Sea has resulted in its rapid expansion to Norway aided by its planktonic larval stages carried by currents and an ability to travel great distances by walking. With a better knowledge of the behavior of organisms during their life-history stages and of the prominent physical vectors from a point of release, theoretical models of dispersal should be possible. Such models would also be valuable tools for predicting the plumes arising from discharges of ships’ ballast water and for evaluating relative risk scenarios of transmitting or receiving exotic species. The reproductive capability of a species is also important. Those that release broods that are confined to the benthos such as the Chinese hat snail Calyptraea chinensis, are likely to remain in one region, and once mature will have a good opportunity for effective reproduction. Species such as Littorina saxatilis have the theoretical capability of establishing themselves from the release of a single female by producing miniature crawlers without a planktonic life stage. In contrast, species with long planktonic stages requiring stable conditions are unlikely to succeed. It is doubtful whether spiny lobsters will become inadvertently transferred in ballast water.

Impacts on Society Exotic species have a wide range of effects: some provide economic opportunities whereas others impose unwanted consequences that can result in serious financial loss and unemployment. In agriculture, the main species utilized in temperate environments for food production were introduced and have taken


340


some thousands of years to develop. In the marine environment, the cultivation of species is relatively new and comparatively few exotic species are utilized. This suggests that in future years an assemblage of exotic species, some presently in cultivation, and others yet to be developed, will form a basis for significant food production. Few exotics in marine cultivation form the basis for subsistence, whereas this is common in freshwater systems. The cultivation of marine species is normally for specialist markets where food quality is an important criterion. Expanding ranges of harmful exotics could erode opportunities by impairing the quality in some way or interfering with production targets. Unfortunately, many shipping port regions are in areas where conditions for cultivation are suitable, either for the practical reason of lower capital costs for management, or for the ease of operation and/or shelter, or since the conditions favor optimal growth and/or the nearby market. The proximity of shipping to aquaculture activities poses the unquantifiable threat that some imported organisms will impair survival, compromise growth, or render a product unmarketable. Diseases of organisms and humans are spreading throughout the world. In the marine environment, a large bulk of biota is in transit in ballast water. Ballasting by ships in port may result in loading untreated discharges of human sewage containing bacteria and viruses that may have consequences for human health once discharged elsewhere. In 1991, at the time of the South American cholera epidemic, caused by Vibrio cholerae, oysters and fish in Mobile Bay, Alabama (USA), were found with the same strain of this infectious bacterium. Ballast water was considered as a possible source of this event. Subsequently five of 19 ships sampled in Gulf of Mexico ports arriving from Latin America were found with this same strain. The epidemic in South America may have been originally sourced from Asia, and may also have been transmitted by ships. Of grave concern is the discovery of new algal toxins and the apparent spread of amnesic shellfish poisoning, diahorritic shellfish poisoning, paralytic shellfish poisoning, and neurological shellfish poisoning throughout the world. The associated algae can form dense blooms that, with onshore winds, can form aerosols that may be carried ashore to influence human health. The apparent increase in the frequency of these events may be due to poor historical knowledge of previous occurrences or to a real expansion of the phenomena. There is good evidence that ballast water may be distributing some of these harmful species. Some of the ‘bloom-forming’ species, such as the naked dinoflagellate Karenia

mikimotoi, are almost certainly introduced and cause sufficiently dense blooms to impair respiration in fishes by congestion of the gills, and can also purge the water column of many zooplankton species and cause mortalities of the benthos. Some introduced invertebrates may act as an intermediate host for human and livestock diseases. The Chinese mitten crab Eriocheir sinensis, apart from being a nuisance species, acts as the second intermediate host for the lung fluke Paragonimus westermanii. The first intermediate stage appears in snails. The Chinese mitten crab has been introduced to the Mediterranean and Black Seas, North America, and northern Europe. Should the lung fluke be introduced, the ability for it to become established now exists where it may cause health problems for mammals, including humans.

Management of Exotic Species Aquaculture

There is an expanding interest in aquaculture as an industry to provide employment, revenue, and food. Already several species in production worldwide contribute to these aims. However, any introduction may be responsible for unwanted and harmful introductions of pests, parasites, and diseases. Those involved in future species introductions should consider the International Council for the Exploration of the Sea’s (ICES) Code of Practice on Introductions and Transfers of Marine Organisms (Table 1). This code takes into account precautionary measures so that unwanted species are unlikely to become unintentionally released. The code also includes provisions for the release of genetically modified organisms (GMOs). These are treated in the same way as if they are exotic species introductions intended for culture. The code is updated from time to time in the light of recent scientific findings. Should this code be ignored the involved parties could be accused of acting inappropriately. By using the code, introductions may take several years before significant production can be achieved. This is because the original broodstock are not released to the wild, only a generation arising from them that is disease-free. This generation must be examined closely for ecological interactions before the species can be freely cultivated. In developing countries it is important that all reasonable precautions are taken to reduce obvious risks. It has been shown historically that direct introductions, even when some precautions have been taken, may lead to problems that can compromise the intended industry or influence other industries, activities, and the environment.



Table 1 Main features of the ICES Code of Practice in relation to an introduction Conduct a desk evaluation well in advance of the introduction, to include: previous known introductions of the species elsewhere; review the known diseases, parasites, and pests in the native environment; understand its physical tolerances and ecological interactions in its native environment; develop a knowledge of its genetics; and provide a justification for the introduction Determine the likely consequences of the introduction and undertake a hazard assessment Introduce the organisms to a secure quarantine facility and treat all wastewater and waste materials effectively Cultivate F1 generation in isolation in quarantine and destroy broodstock Disease-free filial generation may be used in a limited pilot project with a contingency withdrawal plan Development of the species for culture At all stages the advice of the ICES Working Group on the Introductions and Transfers of Marine Organisms is sought. Organisms with deliberately modified heritable traits, such as genetically modified native organisms, are considered as exotic species and are required to follow the ICES Code of Practice.

Experience has shown that good water quality, moderate stocking densities, and meteorological and oceanographic conditions, within the normal limits of species, or the culture system, are of importance for successful cultivation. Maintenance of production in deteriorating conditions, following high sedimentation or pollution, can rarely be achieved by using other introduced species. Exotic species generally used in aquaculture may not always prove to be beneficial. Although the Pacific oyster C. gigas is generally accepted as a useful species, in some parts of the Adriatic Sea it fouls metal ladders, rocks, and stones causing cuts to bathers’ skin, this in a region where revenue from tourism exceeds that from aquaculture. This same species is unwelcome in New South Wales because of competition with the Sydney rock oyster Saccostrea commercialis and there have been attempts to eradicate it. Ecomorphology Organisms can respond to changes in their environment by adapting specific characteristics that provide them with advantage. Sometimes these changes can be noted within a single lifetime, but more usually this takes place over many generations and may ultimately lead to species separation. When considering a species for introduction its morphology may provide clues as to whether it will compete with native species or whether its feeding capabilities or range is likely to

Table 2

341

Treatment measures of ballast water

Disinfection: Tank wall coatings, biocides, ozone, raised temperature, electrical charges and microwaves, deoxygenation, filtration, ultraviolet light, ultrasonification, mechanical agitation, exchanges with different salinity By management: Special shore facilities or lighters (transfer vessels) for discharges, provision of clean water (fresh water) by port authorities, no ballasting when organisms are abundant (i.e., during algal blooms, at night) or of turbid water (i.e., during dredging, in shallows), removal of sediments and disposal ashore. Specific port management plans taking account of local port conditions and seasonality of the port as a donor area Passive effects: Increase time to deballasting, long voyages, reballast at sea.

be distinct and separate, overlapping or coinciding. However, it is not possible to evaluate the overall impacts of a species for introduction in advance of the introduction using morphological features alone. Ships’ Ballast Water

Sterilization techniques of ballast water are difficult, and most ideas are not cost-effective or practical, either because the great volumes of water require large amounts of chemicals or because of the added corrosion to tanks or because the treated water, when discharged, has now become an environmental hazard. Reballasting at sea is the current requirement by the International Maritime Organization (IMO), the United Nations body which deals with shipping. Ballast tanks cannot be completely drained and so three exchanges are required to remove >95% of the original ballast water. However, it is not possible for ships to reballast in mid-ocean in every case. Ships that deballast in bad weather can be structurally compromised or become unstable; this could lead to the loss of the vessel and its crew. A further method under consideration that does not compromise the safety of the vessel is the continuous flushing of water while in passage. Several further techniques have either been considered, researched, or are in development (Table 2). Exotic species management in ballast water is likely to become a major research area into the twenty-first century. Ships’ Hulls

The use of the toxic yet effective organotins as antifouling agents is likely to become phased out over the first decade of the twenty-first century. Replacement coatings will need to be as, or more, effective, if ships are going to manage fuel costs at current levels. Nontoxic coatings or paint coatings


342


containing deterrents to settling organisms are likely to evolve, rather than coatings containing biocides. However, the effectiveness of these coatings will need to take account the normal interdocking times for ships. In some cases robots may be required for reactivating coat surfaces and removing undue fouling. Locations and/or special management procedures where these activities take place need to be carefully planned to avoid establishment of species from the ‘rain’ of detritus from cleaning operations.

Biocontrol Biological control is the release of an organism that will consume or attack a pest species resulting in a population decrease to a level where it is no longer considered a pest. Although there are many effective examples of biological control in terrestrial systems, this has not been practiced in the marine environment using exotic species. However, biological control has been considered in a number of cases. 1. The green alga C. taxifolia has become invasive in the Mediterranean following its likely release from an aquarium; the species presently ranges from the Adriatic Sea to the Balearic Islands and forms meadows over rock, gravels, and sands, displacing many local communities. The introduction of a Caribbean saccoglossan sea slug that does not have a planktonic stage and avidly feeds on this alga has been under consideration for release. These sea slugs were cultured in southern France, but their release to the wild has not been approved. 2. The comb jelly M. leidyi became abundant in the Black Sea in the mid-1980s. It readily feeds on larval fishes and stocks of anchovy and pilchard declined in concert with its expansion. The introduction of either cod Gadus morhua from the Baltic Sea or chum salmon Onchorhynchus gorbusha from North America was considered. Also considered for control was a related predatory comb jelly. However, the predatory comb jelly Beroe became introduced to the Black Sea, possibly in ballast water, and the M. leidyi abundance has since declined. More recently, M. leidyi has entered the Caspian Sea and is repeating the same pattern. 3. The European green crab C. maenas has been introduced to South Africa, Western Australia, and Tasmania as well as to the Pacific and Atlantic coasts of North America. In Tasmania, it avidly feeds on shellfish in culture and on wild mollusks. Here the introduction of a rhizocephalan Sacculina carcini, commonly found within the crab’s

home range and which reduces reproductive output in populations, was considered. However, because there was evidence that this parasite was not host-specific and may infect other crab species, the project did not proceed. 4. The North Pacific sea star A. amurensis has become abundant in eastern Tasmania and may have been introduced there as a result of either ships’ fouling or ballast water releases taken from Japan. This species feeds on a wide range of benthic organisms, including cultured shellfish. A Japanese ciliate Orchitophyra sp. that castrates its host was considered. Exotic species used in biocontrol need to be species-specific. Generalist predators, parasites, and diseases should be avoided. Complete eradication of a pest species may not be possible or cost-effective except under very special circumstances. In order to evaluate the potential input of the control organism, a good knowledge of the biological system is needed in order to avoid predictable effects that may result in a cascade of changes through the trophic system. Native equivalent species of the biological control organism should be sought for first when introducing a biocontrol species. The ICES Code of Practice should be considered and an external panel of consultants should be involved in all discussions. In some cases, control may be possible by developing a fishery for the pest species as has happened in Turkey following the introduction of the rapa whelk R. venosa to the Black Sea and in Norway to control the population of the red king crab P. camtschaticus.

Management Shipping is seen as the most widespread means of disseminating species worldwide and the IMO has held conventions that relate to improved management of ships’ ballast water and sediments and hullfouling management. Aquaculture and the trade in living products also has been responsible for many species transmissions. To this end, ICES has developed a Code of Practice, providing wise advice on deliberate movements of marine species that should be consulted. Species may be released in ignorance to the wild with potential consequences for native communities, such as what may have happened with the introduction of lionfish Pterois volitans to the western North Atlantic. Elimination of an impacting species can rarely be achieved unless found soon after arrival. Regular surveys as well as public participation and awareness may reveal an early arrival. However, this involves up-to-date information. The development of rapid assessment surveys for known



target species may greatly reduce the duration of surveys and provide early information. Such approaches work. The byssate bivalve Mytilopsis sallei was found during a survey in Darwin docks, Australia, and the Mediterranean form of the marine alga C. taxifolia was found in a small bay in California, USA; it was first recognized by a member of the public as the result of an awareness campaign. Such management requires up-to-date dissemination of information to convey current affairs to all stakeholders. Rapid dissemination by free-access online journals, such as Aquatic Invaders, as well as regular scientific meetings, for example, the International Aquatic Invasive Species Conferences, greatly aid in distributing current knowledge. Management of impacting species may take place in different ways according to their mode of life; it will also depend on the vectors involved in their spread. Where natural vectors disperse a species rapidly, controls may be futile. Programs such as the European Union’s specialist projects Assessing Large Scale Risks for Biodiversity with Tested Methods (ALARM) and Delivery Alien Invasive Species Inventories for Europe (DAISIE) to map and manage invading species are likely to lead to new management advances in understanding vector processes and through the development of risk assessments. It is certain that further exotic species will spread and that some will have unpredictable consequences.

Conclusions

343

are not released to the wild in areas that lie beyond their normal range. More effective and less toxic antifouling agents are needed to replace the effective but highly toxic organotin paint applications on ships. This may lead to less toxic port regions which in turn may now become more suitable for invasive species to become established. Despite widespread usage of new antifouling agents on ships’ hulls, unpainted or worn or damaged regions of the hull are areas where fouling organisms will continue to colonize. Ships’ ballast water and its sediments pose a serious threat of transmitting harmful organisms. Future designs of ballast tanks that facilitate complete exchanges, with reduced sediment loading and options for sterilization, could greatly reduce the volume of distributed biota. Exotic species are becoming established at an apparently increasing rate. Some of these will have serious implications for human health, industry, and the environment.

See also Antifouling Materials. Diversity of Marine Species. Habitat Modification. International Organizations. Large Marine Ecosystems. Mariculture of Aquarium Fishes. Pelagic Biogeography. Phytoplankton Blooms.

Further Reading

Exotic species are an important component of human economic affairs and when used in culture or for sport fisheries, etc., require careful management at the time of introduction. They should pass through a quarantine procedure to reduce transmission of any pests, parasites, and diseases. Aquaculture activities, where practicable, should be sited away from port regions, because there is a risk that shipping may introduce unwanted organisms that may compromise aquaculture production. Movements of aquarium species also need careful attention and should be sold together with advice not to release these to the wild. Transported fish are normally stressed and many of these have been shown to carry pathogenic bacteria and parasites. There is a risk that serious diseases of fishes could be transmitted outside of the Indo-Pacific region by the aquarium trade. Trading networks need to consider ways in which organisms transferred alive or organisms and disease agents that may be transferred in or with products,

Cohen AN and Carlton JT (1995) Nonindigenous Aquatic Species in a United States Estuary: A Case Study of the Biological Invasions of the San Francisco Bay and Delta, 246pp. Washington, DC: United States Fish and Wildlife Service. Davenport J and Davenport JL (2006) Environmental Pollution, Vol. 10: The Ecology of Transportation: Managing Mobility for the Environment, 392pp. Dordrecht: Springer (ISBN 1-4020-4503-4). Drake LA, Choi KH, Ruiz GM, and Dobbs FC (2001) Global redistribution of bacterioplankton and viroplankton communities. Biological Invasions 3: 193--199. Hewitt CL (2002) The distribution and diversity of tropical Australian marine bio-invasions. Pacific Science 56(2): 213--222. ICES (2005) Vector pathways and the spread of exotic species in the sea. ICES Co-Operative Report No. 271, 25pp. Copenhagen: ICES. Leppa¨koski E, Gollasch S, and Olenin S (2002) Invasive Aquatic Species of Europe: Distribution, Impacts and Management, 583pp. Dordrecht: Kluwer Academic Publishers (ISBN 1-4020-0837-6).


344


Minchin D and Gollasch S (2003) Fouling and ships’ hulls: How changing circumstances and spawning events may result in the spread of exotic species. Biofouling 19(supplement): 111--122. Padilla DK and Williams SL (2004) Beyond ballast water: Aquarium and ornamental trades as sources of invasive species in aquatic systems. Frontiers in Ecology and Environment 2(3): 131--138. Ruiz GM, Carlton JT, Grozholtz ED, and Hunes AH (1997) Global invasions of marine and estuarine habitats by non-indigenous species: Mechanisms, extent and consequences. American Zoologist 37: 621--632. Williamson AT, Bax NJ, Gonzalez E, and Geeves W (eds.) (2002) Development of a regional risk management framework for APEC economies for use in the control and prevention of introduced pests. APEC MRC-WG Final Report: Control and Prevention of Introduced

Marine Pests. Singapore: Asia-Pacific Economic Cooperation Secretariat.

Relevant Websites http://www.alarmproject.net – ALARM (Assessing Large Scale Risks for Biodiversity with Tested Methods). http://www.aquaticinvasions.ru – Aquatic Invasions (online journal). http://www.daisie.se – Delivering Alien Invasive Species Inventories for Europe (DAISIE). http://www.ices.dk – International Council for the Exploration of the Sea. http://www.imo.org – International Maritime Organization.


EXPENDABLE SENSORS J. Scott, DERA Winfrith, Dorchester, Dorset, UK Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 889–895, & 2001, Elsevier Ltd.

Introduction Expendable sensors represent an approach to ocean measurement in which some degree of measurement precision may be sacrificed in the interests of lower costs and operational expediency. Two requirements of physical oceanography have driven their development: the problem of achieving adequate spatial sampling of the ocean on timescales commensurate with temporal variability; and the requirement by naval forces for under-way assessments of sonar propagation conditions – the first (and still the dominant) application of operational oceanography. The naval requirement first arose in the area of physical oceanography, in the need to know the depth variation of water temperature. In practice, of the three parameters that determine sound speed – temperature, salinity, and pressure – it is temperature that predominates. Pressure is normally deducible with adequate precision from depth, and salinity is normally sufficiently constant to be neglected or simply ‘modeled’ using an archived (T,S) relation. However, salinity may be important near ice, in fiords, and estuaries, and in regions of freshwater influence (ROFIs). The naval requirement is normally for the vertical sound speed profile, and it is the shape of this profile that is important, rather than its mean value. The expendable measurement facility was quickly taken up by the civilian oceanographic community. It gives a means of tackling the problem of how to make synoptic ocean structure measurements where features are likely to move significantly during a survey. A survey with spatial scales small enough to capture interesting features is seriously degraded by their movement and development. Particularly at mid- to high latitudes, a survey using conventional profiling instruments – such as the conductivitytemperature-depth (CTD) probe – cannot be carried out in a time that is small compared with the timescales of motion and development of features such as frontal boundaries and eddies. Surveys are severely limited by deployments that require a vessel to be regularly stationary for casts with a profiling speed of B1 m s 1. Expendable probes allow use at ship speeds up to 20–30 knots,

and air-dropped expendables can clearly outstrip even this. The technique also provides standard results, and the expendable bathythermograph, or XBT, is now considered a central component of global climate monitoring programs such as the Global Ocean Observation System (GOOS). Near-real-time transfer of these data from ships under way is now also an important input to global meteorological forecasting. The XBT was originally intended to improve on the (nonexpendable) mechanical bathythermograph (MBT) which was the principal operational naval device up to the mid-1970s. The MBT was lowered into the water from a vessel, and it inscribed a temperature-depth trace, with a sharp stylus on a small coated glass slide. The temperature-sensing element was a xylene-filled copper tube, whose (temperaturedependent) pressure moved the stylus across the slide via a Bourdon tube. Stylus movement along the slide was determined by a copper bellows, compressed by the increasing water pressure. Data were read from the trace using an optical projector and scale. The XBT was a major advance, allowing operation while under way and dispensing with the intricate measurement routine of the MBT, with the need for a deployment/recovery winch and with the need for calibration. It uses a pre-calibrated thermistor measurement, read onboard in real time. Inference of its depth uses knowledge of its rate of fall through the water. There are now several manufacturers, although the originators. Sippican Inc. (Marion, MA, USA), still lead in the number of available probe types. Current expendable probe capabilities include, in addition to temperature, the measurement of sound speed, conductivity, ocean current, optical properties, and (recently) seabed properties. This review of expendables summarizes the variety of measurable parameters and then outlines the available deployment options. A number of examples are then given of their use in oceanographic research. Expendables specific to naval activities, such as noise-measuring (sonobuoy) systems, will not be covered here, although in some cases they have a limited ocean measurement capability. Air-deployed drifting systems are also omitted.

Expendable Sensor Types Expendable Bathythermograph (XBT)

The purpose of the XBT is to provide a vertical profile of temperature, from the surface to as great a


345

346

EXPENDABLE SENSORS

depth as required, if possible to the seabed. A thermistor measures temperature as the probe descends, and the depth for each measurement is deduced from time of descent using an empirical equation. A number of variants are available with different depth capabilities, the deepest (the T5) reaching 1830 m, which may necessitate data extrapolation in deeper water. Apart from expense, which increases with depth capability, operational constraints become more restrictive for the deeper probes. Whereas the T-7 (760 m) can operate at platform speeds up to 15 knots, a T-5 is limited to 6 knots.1,2Other probe types are (or have been) available,3 but T-7 and T-5 types are in most regular use. The stated accuracy of all probe types is 7 0.151C and 72% of indicated depth, with a depth resolution of 0.65 m. Operational effects of finite depth Various approaches have been adopted to overcome the limited depth capability of XBTs. The best means of doing this is generally accepted to be extrapolation with the help of relevant (same survey) full-depth CTD casts. This has the added advantage of allowing a check on the XBT depth data. In naval operational terms, however,4 this is rarely an option. In conditions where the measured temperature has stabilized at the maximum depth, extrapolation to the seabed using the data trend may be reasonable. A second approach uses extrapolation using archived data, although if these are mismatched, this may be a problem.

Probe design The standard XBT has two main parts: a protective ‘shell’ which remains on the vessel after launch, and the probe itself, which falls through the water and passes data along the

1 The 450m T-4 (460 m, 30 knots), with a rated ship speed of 30 knots, used to be the routine choice for operational use, but this generally gave way to the deeper T-7 at the end of the 1980s. 2 The specified maximum may be exceeded, but prematurebreakage of the wire will then limit the depth of data collected. 3 The T-7, T-5, and T-4 are complemented by the T-6 (460 m, 15 knots), T-10 (200 m, 10 knots), and T-11 (460 m, 6 knots), the latter giving a 0.18 m vertical resolution). A T-7 variation called ‘Deep Blue’ (760 m, 20 knots) was developed for (faster moving) Volunteer Observing Ships. 4 The naval application differs significantly from purely oceanographic applications. At great depth the sound speed increases slowly with depth (pressure), providing weak upward refraction of sound and decreased seabed interaction. Even small temperature gradients may negate this effect, and small errors in the extrapolation may have disproportionate effects.

connecting wire. Electrical contact with the probe is achieved when the unit is loaded into the launcher, allowing initialization of the onboard electronics before the probe is released by withdrawal of a ‘firing pin’ from its tail section. Throughout the operation of the device the probe remains connected to its shell, and thence to the onboard electronics, by two-strand wire. This wire is arranged in two coils, one within the shell, dispensing wire horizontally as the vessel moves away from the launch location, and one within the probe body, which dispenses wire upwards as the probe falls. Data are collected until wire breakage, either when the probe reaches the seabed or (in deeper water) when the wire has been expended without the bottom being reached. If the deploying vessel travels faster than the design speed, the upper coil of wire may be exhausted first. The success of XBT is the result of a number of critical design features. One of these is the small compartment in the shell that contains the electrical contacts, which is designed to avoid the problem of making good instant electrical contacts between a probe, which may have spent many months awaiting use, and a launcher normally sited on an exposed ship deck. In the compartment, the probe contacts are embedded within a thick gelatinous insulator, which is penetrated by the pointed launcher contacts as the breech closes. The material cleans the launcher contacts as they penetrate, and maintains their clean state by the practice of leaving the spent shell in the launcher between probe launches. The free-falling probe itself involves three principal components: the thermistor element, making the temperature measurement; the two-strand wire that connects the thermistor circuit to the onboard electronics; and a weighted, hydrodynamically shaped body. Each of these components plays a vital part in the remarkable success of the instrument as a whole. The thermistor, of course, is indispensable, this small fragment of temperaturesensitive semiconductor providing the measurement capability of the unit. This is positioned in an aperture at the probe tip. The connecting wire may represent a technical achievement at least as great as any of the other XBT components. Two thin strands of copper wire are covered by a thin insulating lacquer which binds them securely but which is sufficiently non-sticky to avoid the problem of self-attachment within the coils, in which hundreds of meters are compactly wound. The probe body consists of a weighty metal nose cone attached to a lightweight faired hollow plastic tail, equipped with fins to ensure vertical travel,


EXPENDABLE SENSORS

and an even metering of wire from the contained coil. The metal cone acts as a ‘sea electrode’ to complete the measurement circuit with the (effectively grounded) vessel. The final part of the instrument is a soft plastic cap, removed before use, which restricts the movement (and possible damage) of the probe in its packaged state. Probe operation Operation of the XBT involves five stages (1) removal of the shell of the previous probe from the launcher; (2) unpacking and insertion of the fresh probe into the launcher, completing the electrical circuit by closing the breech; (3) initialization of the onboard electronics to recognize the probe and to prepare for data acquisition; (4) launch, by withdrawing the ‘firing pin’; and (5) following data acquisition, completion of the data file closure procedure. Data acquisition begins when the probe completes the earth–loop circuit on reaching the water, and continues until the wire breaks. Acquisition may be ended before this if, for example, the probe is known to have reached the seabed. This process is common to all ship-launched probes. Expendable Sound Velocity Probe (XSV)

In the XSV the active sensor, instead of the XBT’s thermistor, is a small sound speed sensor using the ‘sing-around’ principle. In this, an ultrasonic transmitter/receiver pair are arranged with a fixed separation in an electrical circuit with strong feedback. The circuit’s ‘resonance’ (‘sing-around’) frequency is determined mainly by the time acoustic path, and its measurement allows inference of the sound speed. XSVs are almost solely limited to military use,5 principally by operational submarines, and both airlaunched and submarine-launched variants are available for this reason. For naval operations, they offer an improvement over the XBT in regions such as Arctic, Mediterranean, and coastal waters where salinity variations may cause significant sound speed changes. The probes have a specified precision of 70.25 m s 1, and they are available with two main depth options: the XSV-01 (850 m, 15 knots) and the XSV02 (2000 m, 8 knots). Both give depth resolution of 0.32 m. A higher resolution, slower-falling XSV-03 (850 m, 5 knots) is available, giving 0.1 m depth

347

resolution. Depth precision is quoted as 74.6 m or 72% of indicated depth, whichever is greater. Expendable Conductivity–Temperature–Depth Probe (XCTD)

The XCTD is one of the newer expendable probes, measuring conductivity as well as temperature. This is not a simple extension of capability, since precision conductivity measurement is acknowledged to be particularly difficult, and first-time operation must be assured, even following a lengthy unattended shelf-life. A four-electrode configuration is used in a resistive measurement of conductivity, with a thermistor for the temperature measurement. As with other probes, depth is inferred from the fall rate of the probe. In the XCTD the measuring system converts basic C, T data into frequency-modulated signals for transmission along the standard two-wire connection. Each unit is calibrated in a three-saltwater-bath procedure following manufacture, with the resulting calibration coefficients stored in nonvolatile memory in the probe canister. Expendable Current Profiler (XCP)

The Sippican XCP is the first expendable technique for ocean current measurement. Its measurement of current velocity utilizes measurements of the weak6 electrical current generated by the motion of conducting sea water through the Earth’s magnetic field, which is directly proportional to the velocity at any given depth. The falling probe interrupts this current and measures the electrical potential thus produced. The probe spins at a prescribed rate, converting the potential seen by separated electrodes into a sinusoidal signal whose orientation is deduced from a corotating ‘compass’ coil. This allows resolution of the measured current into north and east components; the third measurement made is that of temperature, by the thermistor mounted in the usual probe nose position. The data are passed along the wire as three frequency-modulated signals. The XCP is available only as a stand-alone instrument using a radiofrequency (rf) link, similar to that used for air-launched versions of the XBT and XSV. It may therefore be used from either ship or aircraft, and reception need not be from the deploying platform. An aircraft allows greater reception range, particularly in high sea states. A time interval is allowed between launch and probe release, to allow a deploying ship to move away, reducing its

5

An important exception is their potential use in precision hydrographic surveying, to assess the mean sound speed of the water column.

6 In the measurement, a water velocity of 1 m s about 5 mV at mid-latitudes.


1

corresponds to

348

EXPENDABLE SENSORS

electromagnetic disturbance. Operation involves only the removal of a number of a few items of protective packaging before the unit is dropped into the sea. Energizing of the seawater battery, deployment of the rf antenna, operation of the probe, and eventual scuttling all take place without intervention. The probes have a 1500 m capability, and the 16 Hz (rotation rate determined) sampling rate allows 0.3 m vertical resolution, with a specified velocity resolution of 710 mm s 1 rms and 73% rms horizontal shear current accuracy. The specified temperature resolution is 70.21C.

Expendable Optical Irradiance Probe (XKT)

The XKT, another relatively recent innovation, is used to measure the vertical profile of light penetrating the upper layer of the ocean, allowing an estimate of the optical diffuse attenuation, K, in addition to temperature. The upward-looking probe has a cosine spatial response and is sensitive to the wavelength band 490710 nm, operating down to 200 m with about 0.15 m s 1 vertical resolution. Irradiance is measured over a dynamic range of 105 within 5% log conformity and 106 within 10%. An air-dropped version is available. A second variety of optical properties probe, the XOTD/AXOTD, measuring suspended particle concentrations in addition to temperature, has been reported as ‘under development’. This operates using the scattering of light from an included source and is intended for the particle concentration range 5 mg l 1 to 3 g l 1 in the depth range to 500 m.

Expendable Bottom Penetrometer Probe (XBP)

The XBP is the most recent addition to the armoury of expendables, and it is currently still at the development stage, available by special arrangement for evaluation. The requirement which it fulfils is, once again, driven by military operations, and relates to the need to know certain seabed properties, particularly in shallow water (o200 m). Sonar behavior in shallow water is often determined by the geoacoustic properties of the seabed, and aspects of mine counter-measures, particularly relating to the probability of mine burial, are sensitively dependent on the properties of seafloor sediments. The sensor carried by the XBP in place of the XBT’s thermistor is an accelerometer, whose purpose is to monitor the deceleration of the probe on impact with the seabed. Hard or rocky seabeds involve rapid deceleration, whilst sediments allow a smoother, longer period of retardation.

Normal (Surface Ship) Deployment The standard deployment of all probes is from a surface vessel, normally from the stern, so that the wire emerges freely behind the vessel as it moves away from the launch point. Two types of launcher are normally available: a robust (heavy) military standard deck launcher, usually mounted permanently on deck, and a much smaller hand-held unit which may be used even from small craft. Care is required with either type in strong wind conditions, as the thin wire may be ‘caught’ by the wind and may become entangled with parts of the vessel’s superstructure. (Through-hull launchers are an option used by some military vessels, allowing operation in adverse weather conditions.) Since vessel heading is normally set by survey or operational demands, rather than by wind direction, it is often found that having a pair of launchers, mounted on the port and starboard quarters of the vessel, allows greater flexibility in cross-winds. A hand-held launcher is sometimes used to complement a single deck-mounted launcher for this reason. A less conventional way of addressing this problem is to apply mild restraint to the emerging wire – such as loosely guiding the wire through the fingers. This can reduce the effect of the wind on pulling excess wire from the dispenser within the shell, although care must be taken not to impede its normal flow. Successful use of expendables may also be limited if the vessel is towing equipment, since the XBT wire streaming aft can become fouled by tow cables. The only processing of the data normally needed involves the removal of values from the top 3–5 m, influenced by transient effects as the probe adjusts to the water temperature, and the removal of values obtained after the probe has reached the seabed, an event frequently denoted by a sharp spike in the data.

Deployment Variations The XBT and XSV are available with additional deployment options – air-launched and submarinelaunched, – developed mainly for military use. As was noted above, the XCP is suited to both ship and airborne deployment. The Air-Launched XBT and XSV (AXBT, AXSV)

The measurement parts of the AXBT and AXSV are substantially the same as for the standard probes, including the sensor, the wire attaching the falling probe to the surface unit, and the weighted probe body. In this case, however, the only wire coil used is that within the probe, connecting the sensor to a


EXPENDABLE SENSORS

buoyant electronics package which conditions the temperature signal and communicates it via an rf transmitter link to the launch aircraft. AXBTs and AXSVs are packaged in the standard-size canister used for sonobuoys – the ‘A-size’, 914 mm 124 mm – and they may be launched at air speeds up to 370 knots and altitudes up to more than 9000 feet. Descent is controlled by a parachute, deployed when the buoy leaves the aircraft, and operation begins after a short delay in which the probe reaches temperature stability and the seawater battery (for the rf transmitter) becomes energized. Each unit has a user-selectable rf frequency, which allows simultaneous monitoring of a number of probes. Although as many as 99 channels may be available, the number of probes being deployed is also limited by the number of channels available for simultaneous monitoring by the aircraft. Although transmissions cease when the probe has reached its maximum depth, and the scuttling mechanism is initiated, another probe using the same frequency cannot be launched until this has occurred. The receiver used for AXBTs is a standard unit normally fitted only to military (Maritime Patrol) aircraft. Two types of AXBT are available, designed for maximum depths of 302 m and 760 m. Their spatial resolution is rather better than that of the standard XBT, at about 0.15 m. Specified depth accuracy is 2% of indicated depth, and temperature accuracy is 70.181C. The standard AXSV operators to 760 m, with vertical resolution of about 0.15 m, 2% accuracy, and a specified sound speed accuracy of 70.25 m s 1. In practice, although AXBTs allow rapid deployment over substantial horizontal scales, it is difficult to simultaneously achieve high spatial sampling, because of the combined effect of the finite reload time and the finite number of available receiving channels. To achieve a probe spacing smaller than about 30 km along a single track it is normally necessary to make at least two passes along the track, the second interleaving dropping probes between the stations covered on the first. Despite the operational difficulties, and the relative inaccessibility of such activities to nonmilitary agencies, AXBTs are the only means of executing large-area surveys (hundreds of kilometers) of dynamic regions for which the synoptic requirement requires a time-spread of o1 day. The Submarine-Launched XBT and ASV (SSXBT and SSXSV)

These variants of the XBT and XSV satisfy the uniquely military requirement for a submerged

349

submarine to assess the sonar propagation characteristics of its environment. Without it, a submarine would need either to surface for a conventional deployment or to move vertically, making measurements with onboard sensors. The technique is related to that of the AXBT, in that the temperature profile is measured by a probe falling from a buoyant package which floats at the surface. In this case, however, the package rises to the surface under its own buoyancy following its submerged launch from a signal ejector probe, and remains connected by wire to the submarine. As might be imagined from its requirement to pass probes through the pressure hull of a submarine, this technique involves expensive technology, and is unlikely to be used for purely oceanographic, as opposed to operational, purposes. An equivalent version of the XCTD is understood to be imminent.

Data Recording and Handling Current practice for the recording of data normally involves a PC, with a dedicated electronic interface unit which checks the continuity and integrity of the individual probe electronics before launch and then transfers data only after it has detected the probe reaching the water. Data display options exist using either dedicated display programs or standard data handling routines. As in many oceanographic applications, PCs have dramatically simplified the data recording process. In the years between the emergence of the XBT and the eventual prominence of the PC there was a period in which a dedicated inboard electronic unit was necessary for acquisition and data display using a paper trace. Digital recording on magnetic media (long-since obsolete) was also an option, although the practice of digitization from the paper trace persisted operationally into the 1980s. This inboard equipment frequently tended to be temperamental – a feature that is difficult to forgive when linked to the use of expendable probes. The operational context of XBT data (for meteorological and military purposes) has led to the establishment, by the World Meteorological Organization (WMO) of a standard data exchange format, known as the JJYY format (formerly JJXX the format was officially changed from JJXX to JJYY by the WMO on November 8, 1995), following the alphabetical code used to prefix each ‘bathy report’. Details such as date, time, location, probe type, and recorder type are included in these reports. This format involves reduction of the XBT data to a small number of ‘inflection points’, or ‘break points’,


350

EXPENDABLE SENSORS

which are intended to capture the main features of the temperature profile whilst minimizing the data transfer requirement.

secondary standard before use. Although this method can realistically assess only one temperature, giving an error to be applied as an offset to the subsequent launch, it is a reasonably practical means of improving confidence in the data.

Measurement Precision Assessment of the precision of an expendable device is necessarily limited by the expected loss of the probe following use. However, practical experience normally indicates performance within specified tolerances for the measured variables. Experience is considerably greater for the XBT than for the other probe types, as this is used much more widely for routine purposes. The growing importance of XBTs in the population of global databases has led to some detailed consideration of their precision in the scientific literature. This has centered principally on the derivation of probe depth using the fall-rate equations provided by the manufacturers. Direct verification for individual probe types is not possible because of the large vertical distances involved, and the inappropriateness of attaching verification sensors. The best available means of checking the depth data of expendables, of any kind, is to carry out parallel measurements with a temperature-measuring instrument, such as a CTD, which has a direct pressuremeasuring capability. Distinctive individual temperature features in the water column may then be used to indicate a depth correction for the region of the survey. A number of surveys, several of them involving dense sampling of XBT and CTD casts, have found significant systematic errors greater than the quoted tolerances, the manufacturer’s equation always underestimating the fall rate. A number of alternative fall-rate equations have been proposed, and it appears that a consensus is steadily being reached on the optimum equation. The calibration issue is particularly significant for assessing trends in climate change, and it is acknowledged to be particularly important to follow an agreed standard procedure for handling the known depth errors in databases. Use of a variety of fall-rate equations would lead to major confusion in interpretation of ensembles of data. It is now accepted that only data using the manufacturer’s equation should be used for archived data, and that it should be left to subsequent analysis to make whatever adjustments are felt necessary. It is generally found that temperature errors are within the bounds indicated by the manufacturer. However, it has been indicated that performance may be improved by calibrating each probe against a

Use for Ocean Surveying Although XBT use for purely oceanographic survey purposes is probably still dominated by naval requirements this is an effective way of enhancing a traditional survey using CTD casts. Without impact on survey time it is possible to increase the spatial sampling rate by a factor of two or more by interpolating XBT launches between CTDs. Another common use is in the support of surveys undertaken with towed undulators or instrument chains, XBTs allowing a degree of downward extrapolation of the detailed upper-layer data that these collect. A third common use of XBTs is as a ‘fall-back’ option for use when weather conditions are unsuitable for other equipment. AXBTs, and their deployment, are considerably more expensive, and tend to be used only when rapidity is essential. For example, the use of repeated large-scale AXBT surveys of the highly dynamic Iceland-Faroes frontal zone has been reported. AXBTs have also been used to give near-synoptic sections using aircraft underflights of satellite altimeter tracks, to validate the use of residual height data for ocean monitoring. The viability of these techniques depends on the proportion of budget available for expendable items, and the expense of the probes must be seen in the context of the high basic cost of trials. The other main contribution made to ocean science by expendables relates to deductions made from data archives. These are frequently dominated by XBT data and they often allow spatial and temporal coverage of regions that would otherwise have insufficient coverage for reliable deductions to be made. None of the other probe types described here comes close to the XBT in its contribution to the science. Perhaps the most detailed and intensive use of the other expendables has involved the XCP, whose data (drawn from a number of different ocean regions) have been used to draw conclusions about the horizontal shear environment of the ocean.

See also Acoustics, Deep Ocean. Acoustics, Shallow Water. CTD (Conductivity, Temperature, Depth) Profiler.


EXPENDABLE SENSORS

Data Assimilation in Models. Inherent Optical Properties and Irradiance. Ships. Sonar Systems.

Further Reading Black PG, Elsberry RL, and Shay LK (1988) Airborne surveys of ocean current and temperature perturbations induced by hurricanes. Advances in Underwater Technology, Ocean Science and Offshore Engineering 16: 51--58. Budeús G and Krause G (1993) On-cruise calibration of XBT probes. Deep-Sea Research 40: 1359--1363. Carnes MR and Mitchell JL (1990) Synthetic temperature profiles derived from Geosat altimetry: Comparison with

351

air-dropped expendable bathythermograph profiles. Journal of Geophysical Research 95: 17979--17992. Smart JH (1984) Spatial variability of major frontal systems in the North Atlantic Norwegian Sea area: 1980–1981. Journal of Physical Oceanography 14: 185--192. Smart JH (1988) Comparison of modelled and observed dependence of shear on stratification in the upper ocean. Dynamics of Atmospheres and Oceans 12: 127--142. Stoll RD and Akal T (1999) XBP – Tool for rapid assessment of seabed sediment properties. Sea Technology February: 47--51. Thadathil P, Ghosh AK, and Muraleedharan PM (1998) An evaluation of XBT depth equations for the Indian Ocean. Deep-Sea Research 45: 819--827.


FIORD CIRCULATION hydrodynamic processes and simple quantitative models of the main types of circulation are presented.

A. Stigebrandt, University of Gothenburg, Gothenburg, Sweden & 2009 Elsevier Ltd. All rights reserved.

Basic Concepts in Fiord Descriptions Introduction Fiords are glacially carved oceanic intrusions into land. They are often deep and narrow with a sill in the mouth. Waters from neighboring seas and locally supplied fresh water fill up the fiords, often leading to strong stratification. During transport into and stay in the fiord, mixing processes modify the properties of imported water masses. From the top downward, the fiord water is appropriately partitioned into surface water, intermediary water, and, beneath the sill level, basin water. Fiord circulation is forced both externally and internally. External forcing is provided by temporal variations of both sea level (e.g., tides) and density of the water column outside the fiord mouth. Internal forcing is provided by freshwater supply, winds, and tides in the fiord. The response of circulation and mixing in the different water masses to a certain forcing depends very much on characteristics of the fiord topography. The circulation of basin water is critically dependent on diapycnal mixing. Hence we focus on fiord circulation from a hydrodynamic point of view. Major

Some key elements of the hydrography and dynamics of fiords are shown in Figure 1. The surface water may have reduced salinity due to freshwater supply. It is kept locally well mixed by the wind and the thickness is typically of the order of a few meters. The thickness of the intermediary layer, reaching down to the sill level, depends strongly on the sill depth. This layer may be thin and even missing in fiords with shallow sills. Surface and intermediary waters have free connection with the coastal area through the fiord’s mouth. Basin water, the densest water in the fiord, is trapped behind the sill. It is vertically stratified but the density varies less than in the layers above. A typical vertical distribution of density, r(z), in a strongly stratified fiord is shown in Figure 1. The area (water) outside the fiord is denoted here as coastal area (water). In most fiords, temporal variations of the density of the coastal water are crucial for the water exchange, both above and below the sill level. Surface and intermediary waters in short fiords with relatively wide mouths may be exchanged quickly (i.e., in days). The vertical stratification in the intermediary layer in such a fiord is usually quite similar to the Fresh water

Wind Surface water

Estuarine circulation

Entrainment Initial mixing (z)

Coastal water

Intermediary water

Dense bottom current (intermittent)

ent

ainm

Entr

Sill

al Intern s e v a w

Basin water

Figure 1 Basic features of fiord hydrography and circulation. Turbulent mixing in the basin water occurs mainly close to the bottom boundary, with increased intensity close to topographic features where internal tides are generated.


353

354

FIORD CIRCULATION

stratification outside the fiord although there is some phase lag, with the coastal stratification leading before that in the fiord. Short residence time for water above sill level means that the pelagic ecology may change rapidly due to advection. Denser coastal water, occasionally appearing above sill level, intrudes into the fiord and sinks down along the seabed. It then forms a turbulent dense bottom current that entrains ambient water, decreasing the density and increasing the volume flow of the current (Figure 1). The intruding water replaces residing basin water. During so-called stagnation periods, when the density of the coastal water above sill level is less than the density of the basin water, a pycnocline develops at or just below sill depth in the fiord. In particular, during extended stagnation periods, oxygen deficit and even anoxia may develop. The density of basin water decreases slowly due to turbulent mixing, transporting less-dense water from above into this layer. Tides are usually the main energy source for deep-water turbulence in fiords. Baroclinic wave drag, acting on barotropic tidal flow across sills separating stratified basins, is the process controlling this energy transfer.

Major Hydrodynamic Processes in Fiord Circulation Understanding fiord circulation requires knowledge of some basic physical oceanographic processes such as diapycnal mixing and a variety of strait flows, regulating the flow through the mouth. Natural modes of oscillation (seiches) as well as sill-induced processes like baroclinic wave drag, tidal jets, and hydraulic jumps may constitute part of the fiord response to time-dependent forcing. Flow through fiord mouths is driven mainly by barotropic and baroclinic longitudinal pressure gradients. The barotropic pressure gradient, constant from sea surface to seabed, is due to differing sea levels in the fiord and the coastal area. Baroclinic pressure gradients arise from differing vertical density distributions in the fiord and coastal area. The baroclinic pressure gradient varies with depth. Different types of flow resistance and mechanisms of hydraulic control may modify flow through a mouth. Barotropic forcing usually dominates in shallow fiord mouths while baroclinic forcing may dominate in deeper mouths. Three main mechanisms cause resistance to barotropic flow in straits. First is the friction against the seabed. Second is the large-scale form drag, due to large-scale longitudinal variations of the vertical cross-sectional area of the strait causing contraction followed by expansion of the flow. And third is the

baroclinic wave drag. This is due to generation of baroclinic (internal) waves in the adjacent stratified basins. Barotropic flow Q through a straight, rectangular, narrow, and shallow strait of width B, depth D, and length L, connecting a wide fiord and the wide coastal area may be computed from Q2 ¼

2gDZB2 D2 1 þ 2CD ðL=DÞ

½1

The sign of Q equals that of DZ ¼ ho hi, ho (hi) is the sea level in the coastal area (fiord). This equation contains both large-scale topographic drag and drag due to bottom friction (drag coefficient CD). Largescale topographic drag is the greater of the two in short straits (i.e., where L/Do2/CD). To get first estimates of baroclinic flows and processes, it is often quite relevant to approximate a continuous stratification by a two-layer stratification, with two homogeneous layers of specified thickness and density difference Dr on top of each other. The reduced gravity of the less-dense water is g0 ¼ gDr/r0, where g is the acceleration of gravity and r0 a reference density. In stratified waters, fluctuating barotropic flows (e.g., tides) over sills are subject to baroclinic wave drag. This is important in fiords because much of the power transferred to baroclinic motions apparently ends up in deep-water turbulence. Assuming a twolayer approximation of the fiord stratification, with the pycnocline at sill depth, the barotropic to baroclinic energy transfer Ej from the jth tidal component is given by Ej ¼

r0 2 2 A2f Hb o a ci 2 j j Am Hb þ Ht

½2

Here oj is the frequency and aj the amplitude of the tidal component. Af is the horizontal surface area of the fiord, Am the vertical cross-sectional area of the mouth, Ht (Hb) sill depth (mean depth of the basin water), and sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi Ht Hb c i ¼ g0 Ht þ Hb the speed of long internal waves in the fiord. Baroclinic wave drag occurs if the speed of the barotropic flow in the mouth is less than ci. If the barotropic speed is higher, a tidal jet develops on the lee side of the sill together with a number of flow phenomena like internal hydraulic jumps and associated internal waves of supertidal frequencies.


FIORD CIRCULATION

Baroclinic flows in straits may be influenced by stationary internal waves imposing a baroclinic hydraulic control. For a two-layer approximation of the stratification in the mouth the flow is hydraulically controlled when the following condition, formulated by Stommel and Farmer in 1953, is fulfilled: u21m u22m þ ¼1 g0 H1m g0 H2m

½3

Here u1m (u2m) and H1m (H2m) are speed and thickness, respectively, of the upper (lower) layer in the mouth. Equation [3] may serve as a dynamic boundary condition for fiord circulation as it does in the model of the surface layer presented later in this article. It has also been applied to very large fiords like those in the Bothnian Bay and the Black Sea. Experiments show that superposed barotropic currents just modulate the flow. However, if the barotropic speed is greater than the speed of internal waves in the mouth these are swept away and the baroclinic control cannot be established and the transport capacity of the strait with respect to the two water masses increases. In wide fiords, the rotation of the Earth may limit the width of baroclinic to the order of pcurrents ffiffiffiffiffiffiffiffiffiffi the internal Rossby radius g0 H1 =f . Here H1 is the thickness of the upper layer in the fiord and f the Coriolis parameter. The outflow from the surface layer is then essentially geostrophically balanced and the transport is Q1 ¼ g0 H12 =2f . This expression has been used as boundary condition for the outflow of surface water, for example, from Kattegat and the Arctic Ocean through Fram Strait. Diapycnal (vertical) mixing processes may modify the water masses in fiords. In the surface layer, the wind creates turbulence that homogenizes the surface layer vertically and entrains seawater from below. The rate of entrainment may be described by a vertical velocity, we, defined by we ¼

m0 u3 g0 H1

½4

Here u is the friction velocity in the surface layer, linearly related to the wind speed, and m0 (B0.8) is essentially an efficiency factor, well known from seasonal pycnocline models. H1 is the thickness of the surface layer in the fiord and g0 the buoyancy of surface water relative to the underlying water. Buoyancy fluxes through the sea surface modify the entrainment velocity. Due to their sporadic and ephemeral character, there are only few direct observations of dense bottom currents in fiords. However, from both observations and modeling of dense bottom currents in the

355

Baltic it appears that entrainment velocity may be described by eqn [4]. For this application, u is proportional to the current speed, H1 is the current thickness, and g0 the buoyancy of ambient water relative to the dense bottom current. Observational evidence strongly supports the idea that diapycnal mixing in the basin water of most fiords is driven essentially by tidal energy, released by baroclinic wave drag at sills. The details of the energy cascade, from baroclinic wave drag to smallscale turbulence, are still not properly understood. Figure 1 leaves the impression that energy transfer to small-scale turbulence and diapycnal mixing takes place in the inner reaches of a fiord. Here internal waves, for example, waves generated by baroclinic wave drag at the sill, are supposed to break against sloping bottoms. A tracer experiment in the Oslo Fiord suggests that mixing essentially occurs along the rim of the basin. Recently, the temporal and spatial distributions of turbulent mixing in the basin waters of a few fiords have been mapped. Measurements in the Gullmar Fiord suggest that much of the mixing takes place close to the sill where most of the barotropic to baroclinic energy transfer takes place. In a column of the basin water the mean rate of work against the buoyancy forces is given by W ¼ W0 þ

Rf

Pn

j¼1

At

Ej

½5

Here W0 is the nontidal energy supply, n the number of tidal components, and Ej may be obtained from eqn [2]. At is the horizontal surface area of the fiord at sill level and Rf the flux Richardson number, the efficiency of turbulence with respect to diapycnal mixing. Estimates from numerous fiords show that RfB0.06. Experimental evidence shows that in fiords with a tidal jet at the mouth, most of the released energy dissipates above sill level and only a small fraction contributes to mixing in the deep water.

Simple Quantitative Models of Fiord Circulation The Surface Layer

The upper layers in fiords may be exchanged due to so-called estuarine circulation, caused by the combination of freshwater supply and vertical mixing. This is essentially a baroclinic circulation, driven by density differences between the upper layers in the fiord and coastal area, respectively. However, if the sill is very shallow, the water exchange tends to


356

FIORD CIRCULATION

be performed by barotropic flow with alternating direction, forced by the fluctuating sea level outside the fiord. The following equations describe the steady-state volume and salt conservation of the surface layer: Q1 ¼ Q2 þ Qf

½6

Q 1 S 1 ¼ Q 2 S2

½7

Here Qf is the freshwater supply, Q1 (Q2) the outflow (inflow) of surface (sea-) water, and S1 (S2) the salinity of the surface water (seawater). Equations [6] and [7] give: Q1 ¼ Qf

S2 S1 S2

½8

This equation has been used throughout the past century for diagnostic estimates of the magnitude of estuarine circulation from measurements of S1, S2, and Qf. Density varies with both salinity and temperature. In brackish waters, density (r) variations are often dominated by salinity (S) variations. For simplified analytical models, one may take advantage of this and use the equation of state for brackish water: r ¼ rf ð1 þ bSÞ

½9

Here rf is the density of fresh water and the so-called salt contraction coefficient b equals 0.000 8 S 1. The density difference Dr between two homogeneous layers, with salinity difference DS, then equals rfbDS, and the buoyancy g0 ¼ gDr/r equals gbDS. A continuous stratification in a salt-stratified system may be replaced by a dynamically equivalent two-layer stratification. This requires that the two-layer and observed stratification (1) contain the same amount of fresh water and (2) have the same potential energy. Stationary estuarine circulation in fiords with deep sills To investigate how salinity S1 and thickness H1 of the surface layer depend on wind and freshwater supply, the following simple model may be illustrative. It is assumed that the fiord mouth is deep and even narrow compared to the fiord. Then the so-called compensation current into the fiord is deep and slow compared with the current of outflowing surface water. The hydraulic control condition in the mouth, eqn [3], is then simplified to u21m ¼ g0 H1m . The thickness of the surface layer H1 in the fiord is related to that in the mouth by H1 ¼ jH1m . Entrainment of seawater of salinity S2 into the surface layer is described by eqn [4]. Under

these assumptions, expressions for the thickness H1 and salinity S1 of the surface layer in the fiord may be derived: 2 1=3 Q G þ j 0 f2 H1 ¼ g Bm 2Qf g0

S1 ¼

S2 G 0 2 #1=3 g G þ 2j Q5f Bm "

½10

½11

Here Bm is the width of the control section in the mouth, G ¼ CWs3Af, C ¼ 2.5 10 9 is an empirical constant containing, among others, the drag coefficient for air flow over the sea surface, Ws the wind speed and g0 ¼ gbS2. Theoretically, the value of j is expected to be in the range 1.5–1.7 for fiords where Bm/Bf r 1/4 where Bf is the width of the fiord inside the mouth. The value of j should be smaller for wider mouths. Observations in fiords with narrow mouths give j values in the range 1.5–2.5. The left term in the expression for H1 is the socalled Monin–Obukhov length, known from the theory of geophysical turbulent boundary layers with vertical buoyancy fluxes, and the right term is the freshwater thickness H1f, hydraulically controlled by the mouth. The salinity of the surface layer S1 increases with increasing wind speed and decreasing freshwater supply. For a given freshwater supply, strong winds may apparently multiply the outflow as compared with Qf (cf. eqn [8]). The freshwater volume in the fiord is Vf ¼ H1fAf. The residence time of fresh water in the fiord, tf ¼ Vf /Qf is given by tf ¼ jAf

1 g0 B2m

1=3

1=3

Qf

½12

The residence time thus decreases with the freshwater supply. It should be noted that H1f, tf, and Vf are independent of the rate of wind mixing. Water exchange through very shallow and narrow mouths In very shallow and narrow fiord mouths, barotropic flow usually dominates. The instantaneous flow which can be estimated using eqn [1] is typically unidirectional and the direction depends on the sign of DZ, the sea level difference across the mouth. If the mouth is extremely shallow and narrow, tides and other sea level fluctuations in the fiord will have smaller amplitude than in the coastal area due to the choking effect of the mouth. A number of choked fiords and other semi-enclosed


FIORD CIRCULATION

water bodies have been described in the literature, such as Framvaren and Nordaasvannet (Norway), Sechelt Inlet (Canada), the Baltic Sea, the Black Sea, and tropical lagoons. The Intermediary Layer

Density variations in the coastal water above sill level give rise to water exchange across the mouth, termed intermediary water exchange. The stratification in the fiord strives toward that in the coastal water. Intermediary circulation increases in importance with increasing sill depth. It has been found that baroclinic intermediary circulation is the dominating circulation component in a majority of Scandinavian and Baltic fiords and bays. This is probably true also in other regions, although there have been few investigations quantifying intermediary circulation. Despite its often-dominating contribution to water exchange in fiords, the intermediary circulation has remained astonishingly anonymous and in many studies of inshore waters even completely overlooked. One obvious reason for this is that a simple formula to quantify the mean rate of intermediary water exchange, Qi, has been available only during the last decade (eqn [13]). sffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffiffi gDM Qi ¼ g Bm Ht Af r

below sill level. In most fiords the filling time is much shorter than the stagnation time. The spectral distribution of the density variability in the coastal water determines the recurrence time of water of density higher than a certain value. Knowing the spectral distribution of the variability of the coastal density and the rate of density decrease of the basin water due to diapycnal mixing, one may estimate the recurrence time of water exchange. In fiords with short filling time, this should be inversely proportional to the rate of diapycnal mixing. A very long stagnation period may occur only if the rate of vertical mixing is very low and the density in the coastal water has a long-period component of appreciable amplitude. A rough estimate of the mean rate of diapycnal mixing in the basin water may be obtained from eqn [14]: dr CW ¼ 2 dt gHb

½14

Here the empirical constant C equals 2.0 and W may be obtained from eqn [5]. The vertical diffusivity k at the level z in the basin water may be computed from the empirical expression in eqn [15]:

½13

Here the dimensionless empirical constant g equals 17 10 4, as estimated for Scandinavian conditions, and DM the standard deviation of the weight of the water column down to sill level (kg m 2) in the coastal water. The latter should be a hydrodynamically reasonable measure of the mean strength of the baroclinic forcing. Statistics of scattered historic hydrographic measurements may be used to compute DM. Equation [13] should be regarded as a precursor to a formula, which is yet to be developed, accommodating for the frequency dependence of DM. A conservative estimate of the mean residence time for water above the sill is ti ¼ Vi/Qi, where Vi is the volume of the fiord above sill level. The residence time may be shorter if other types of circulation contribute to the water exchange.

357

kðzÞ ¼

W=Hb NðzÞ 1:5 c k ¯ ¯2 N rN

½15

¯ is the volume-weighted vertical average of Here N the buoyancy frequency N(z) and ck (B1) is an empirical constant. If the filling time of the fiord basin is very long, the basin will be filled not only with the densest but also with less dense coastal water. The basin water in such a fiord will thus have lower density than the basin water in a neighboring fiord with similar conditions except for a much shorter filling time. If the filling time is sufficiently long, this will determine the residence time, which will not change if the rate of vertical mixing changes. The fiord basin is then said to be overmixed. The Baltic and the Black Seas are two overmixed systems for which the transport capacities of the mouths determine the residence time.

The Basin Water

The time between two consecutive exchanges of basin water, often called the residence time, can be partitioned into the stagnation time (when there is no inflow into the basin) and the filling time (the time it takes to fill the basin with new deep water). The filling time is determined by the flow rate of new deep water through the mouth and the fiord volume

Conclusions This article on fiord circulation demonstrates that several oceanographic processes of general occurrence are involved in fiord circulation. Being sheltered from winds and waves, fiords are excellent large-scale laboratories for studying these processes. The mechanics of water exchange of surface and


358

FIORD CIRCULATION

intermediary layers, as described here, should apply equally well to narrow bays lacking sills.

W Ws z b

Nomenclature aj Af Am

amplitude of the jth tidal component horizontal surface area of the fiord vertical cross-sectional area of the mouth horizontal surface area of the fiord at At sill depth B width of the mouth channel width of the fiord inside the mouth Bf width of the mouth at the control Bm section empirical constant (E1) ck C empirical constant ( ¼ 2.5 10 9) CD drag coefficient for flow over the seabed D depth of the mouth channel barotropic to baroclinic energy Ej transfer from the jth tidal component g acceleration of gravity ( ¼ gDr/r0) buoyancy g0 G wind factor sea level in the fiord hi sea level in the coastal area ho H1m(H2m) thickness of upper (lower) layer in the mouth mean depth of the basin water Hb sill depth Ht L length of mouth channel efficiency factor ( ¼ 0.8) m0 N(z) buoyancy frequency at depth z ¯ N volume-weighted buoyancy frequency in the deepwater flow out of the surface layer Q1 flow into the fiord beneath the surface Q2 layer freshwater supply Qf Rf Richardson flux number, efficiency factor salinity of upper (lower) layer S1 (S2) t time u1m (u2m) speed of upper (lower) layer in the mouth friction velocity u volume of fresh water in the upper Vf layer

g DM

DZ Dr r0 rf r(z) tf j oj

rate of work against buoyancy forces in a column of basin water wind speed vertical coordinate salt contraction of water ( ¼ 0.000 8 S 1) empirical constant ( ¼ 17 10 4) standard deviation of weight of water column between sea surface and sill level sea level difference across the mouth (ho – hi) density difference reference density density of fresh water density as function of depth z residence time for fresh water in the upper layer contraction factor ( ¼ 1.5 for fiords with narrow mouths) frequency of the jth tidal component

See also Flows in Straits and Channels. Internal Tides.

Further Reading Arneborg L, Janzen C, Liljebladh B, Rippeth TP, Simpson JH, and Stigebrandt A (2004) Spatial variability of diapycnal mixing and turbulent dissipation rates in a stagnant fiord basin. Journal of Physical Oceanography 34: 1679--1691. Aure J, Molvær J, and Stigebrandt A (1997) Observations of inshore water exchange forced by fluctuating offshore density field. Marine Pollution Bulletin 33: 112--119. Farmer DM and Freeland HJ (1983) The physical oceanography of fiords. Progress in Oceanography 12: 147--220. Freeland HJ, Farmer DM and Levings DC (eds.) (1980) Fiord Oceanography, 715pp. New York: Plenum. Stigebrandt A (1999) Resistance to barotropic tidal flow by baroclinic wave drag. Journal of Physical Oceanography 29: 191--197. Stigebrandt A and Aure J (1989) Vertical mixing in the basin waters of fiords. Journal of Physical Oceanography 19: 917--926.


FIORDIC ECOSYSTEMS K. S. Tande, Norwegian College of Fishery Science, Tromsø, Norway Copyright & 2001 Elsevier Ltd. . This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 902–909, & 2001, Elsevier Ltd.

Introduction Fiords and semienclosed marine systems are characterized by distinct vertical gradients in environmental factors such as salinity, temperature, nutrients, and oxygen; sampling of biological variables is considerably less adversely influenced by currents and weather conditions than in offshore systems. Therefore, the fiords are particularly well-suited for detailed studies of the pelagic habitat and have traditionally attracted marine scientists, mainly as a site for curiosity-driven research. They are easily accessible and have therefore been used as experimental laboratories for testing new methodologies, exploring trophic relations, and building new theories. Compared with the fisheries on the shelf outside or in open oceanic waters, the catch from the fiords does not seem proportionately great. Nevertheless, fiords play an important role in many local communities in productive coastal areas in Norway, Scotland, British Columbia, and Greenland. There are fisheries locally renowned for local herring stocks – Loch Fyne for example – and for early year classes of Atlantic– Scandian herring. The salmon fisheries of British Columbia, Scotland, and Norway, highly seasonal affairs catching the fish as they return from the open sea to spawn in fresh water, are also a prominent resource. In more recent time, fiords have become refuse pits of cities and highly populated areas. A growing awareness of the importance of these sites for the bordering societies has led to a more cautionary use of the fiords. This is in particular true for the development of salmon and mariculture production facilities, where future demand for area is being met in coastal developmental plans developed from our knowledge of the physical and the biological setting of the particular fiord habitats. The trophic relationships play an essential role in shaping the pattern of species interactions. Delineating their existence in the form of food chains and webs provides important knowledge of ecosystem functioning and dynamics. This basic knowledge is essential for the understanding that has to form the basis for ecosystem and management

models for the marine biota; this article presents some of this generic knowledge for fiordic ecosystems.

General Features Globally, most of the fiords are found in high latitudes: in the Northern Hemisphere north of 501N in Canada, the United States, Greenland, Scandinavia, Scotland and Svalbard; and south of 401S in Peru, Chile, and New Zealand. They are subjected to temperate, boreal, or arctic climates, and have many oceanographic processes that occur over a wide range of space and timescales. At the high-frequency end of the spectrum, tidal flows can generate turbulence, internal waves, and eddies, while processes such as renewal of oxygen in deep enclosed basins may in extreme cases have timescales of tens of years. Another property of fiords is the presence of strong gradients in a number of abiotic factors, which may affect organisms. The gradients may exist in the horizontal or vertical plane and may be permanent or variable. Topography and Estuarine Circulation

The fiord systems of the world are all quite young in evolutionary time, yet they offer opportunities that are quite unique among coastal systems. Their common origin as glaciated coastal U-shaped valleys and the common hydrographic conditions give them a number of characteristic features. They are generally long, narrow, and often deeper than the continental shelf outside. A deep basin connects directly or indirectly with the open sea at one end over a relatively shallow sill, typically one-half to one-tenth the basin depth (Figure 1). The sill, formed by moraine material as it accumulated at the ice edge, limits the exchange of water between the fiord and the coastal region outside. Most fiords have large rivers that often enter at the head of the fiord, but a substantial amount of fresh water is drained into the fiord from the bordering mainland. The water masses can be separated into three vertical layers: brackish surface, intermediate, and basin water. The brackish surface layer varies in time and space, dictated by the fresh water runoff, wind, and tide. At the river outlet, a local elevation of the surface generates a density-driven fresh water current out of the fiord, which is gradually being mixed with saline water from below. This vertical mixing process is called entrainment and leads to an overall increase in


359

360

FIORDIC ECOSYSTEMS

River outlet Upper layer Entrainment Intermediate layer f Sill

el

Sh

S

Coastal water

Atlantic water

Basin layer Sill N

Figure 1 Basic topography and water layers of a fiord with estuarine circulation.

salinity in the surface brackish water. Entrainment leads to a compensation current below that is going in the opposite direction. This pattern of water transport is called estuarine circulation and is a prominent feature in most fiords. Owing to the Coriolis force there is a tendency for water currents to be deflected to the right (in the Northern Hemisphere), which will generate a larger outflow of surface water on the right side of the fiords compared to the left side. The tides vary locally, but will in general add only little to the net water transport. Therefore the windgenerated and estuarine circulations, are important for the exchange of water between fiords and the outside shelf. The wind-driven circulation is dictated by the wind stress on the surface and the vertical density gradient. The effect of the wind decreases vertically with increasing density gradient. Since the wind blows in or out the fiord, the wind-driven circulation sets up the estuarine circulation, as either an inward or an outward flowing surface layer depending on the magnitude, duration, and direction of the wind. The main reason for water exchange in the intermediate layer is the existence of density gradients of water between the fiord and the shelf outside, which follow wind driven upwelling and downwelling at the shelf (Figure 2). Characteristically, this exchange occurs mainly as a two-layer circulation with inflow in the surface layer and a compensating outflow at the lower intermediate layer. This happens when the prevailing wind is northward along the western side of the continent (downwelling), and an oppositely directed circulation occurs when the wind is coming from the north (upwelling). The sill creates a natural barrier that prevents water exchange between the basin and the corresponding depth outside. Partial or total renewal of the basin water takes place when water of higher density than the deep basin water is found above the sill. The major periods of deep-water renewals are found in late winter during upwelling conditions on the shelf outside. Total renewal of the deep water in the fiords happens occasionally, with a frequency of 5–10

Coastal Atlantic water water

Figure 2 General pattern of wind-induced water exchanges between coastal and fiords on the western European continental margins, with prevailing northerly wind (upper panel) with exchange of basin water, and southerly wind (lower panel) with exchange of water in the intermediate layer.

years. The basin water in most fiords is being renewed at a rate that prevents anoxia. Nevertheless, in some fiords with very shallow sills, the renewal of the basin water is low and takes place through vertical diffusion. Here one can find anoxic conditions with high levels of hydrogen sulfide, due to the organic drain-off from land and sedimentation from the surface plankton production. The above outline of the basic physical functioning of fiords underlines the open-ended nature of fiord communities, which nevertheless seem to maintain individual and group characteristics in terms of their biology. There may be a useful distinction between fiords sensu stricto and smaller, more shallow enclosed systems (i.e., ‘polls’). The basic distinction between these two systems is that fiords have a sill depth greater than the depth of the pycnocline, while shallow enclosed systems normally have a sill depth less than the depth of the pycnocline. In biological terms, the latter are in general an ultraplankton– microzooplankton community, whereas the fiord community is chiefly a net phytoplankton–macro/ megazooplankton community. Fresh Water Runoff and Nutrient Cycling

The fresh water runoff plays a major modifying role in the dynamics of the lower trophic organisms in a fiord. The magnitude of this fresh water runoff varies depending on the extent of the surrounding landmass. Fiords can be classified as high-runoff and lowrunoff, arbitrarily separated at an annual mean river


FIORDIC ECOSYSTEMS

runoff of 150 m3 s1. On the other hand, peak values of a few weeks’ duration are found in the range from 500 to 20 000 m3 s1. There is a strong seasonal variation in fresh water runoff, with annual maximum around April in southern fiords and around June in subarctic regions. Low runoff in the beginning of the season will enhance the stabilization of the water masses, facilitating the onset of the spring bloom. Later, increased runoff enhances the washout of the surface water layer, and modifies the phytoplankton biomass and species composition. Fresh water runoff plays a minor direct role in the nutrient cycle of fiords, where the nitrate supplied has been found to account for a maximum of 10–20% of the potential uptake by phytoplankton in June in western Norwegian fiords. Indirect factors influencing the transport of new nutrients into the photic zone from below are therefore of major importance for the new production within the fiord. Although most of the nitrate loss out of fiords takes place below the euphotic zone, such losses may affect the future vertical transport of nitrate from the nonphotic to the photic zone. Hence, advective nutrient loss in the layer close to the photic zone may lead to reduced new production within the fiord. Since loss rates of nutrient appears to exceed loss rates from phytoplankton, an advective nutrient exchange, even below the euphotic zone, may therefore be of greater importance to the primary production than the exchange of the phytoplankton itself. The above picture prevails under specific wind conditions on the coast, and in case of reversed winds a nutrient loss may be turned into net nutrient supply to the fiord.

Seasonality in Energy Input to the System Primary Production

Year-round measurements of primary production exist for a number of fiords and demonstrate that they are in the range of the general level of production in coastal waters of 100–150 g C m2 y1 (Table 1). Some of the data indicate that remarkably high levels of primary production can exist where boundary conditions or land runoff enhance the nutrient load to the fiord. Vertical stability can also generate off-seasonal blooms, even in November and December in fiords at lower latitudes, but it is unlikely that these blooms effectively stimulate secondary production. A significant proportion of the annual phytoplankton production occurs before the macrozooplankton grazing population becomes established. The estimated annual flux of phytoplankton based

361

Table 1 Estimates of primary production in representative fiords on the Northern Hemisphere Region

Norway Balsfiorden Lindspolls Sweden Kungsbackafjord Greenland Godthaabfjord Canada How Sound entrance entrance inner stations inner stations Port Moody Arm Indian Arm and the Narrows regions Strait of Georgia River Plume USA Puget Sound Port Valdez Valdez Arm

Year

Primary production (g Cm2 y1)

1977 1976

110 90–100

1970

100

1953–56

98

1973 1974 1973 1974 1975–76 1975–76

300 516 118 163 532 260

1965–68 1975–77

120 149

1966–67 1971–72 1971–72

465 150 200

carbon to the bottom of fiords is B10% of the total particulate material sedimentation. This means the majority of the particulate material reaching the bottoms of the fiord is inorganic. Although there is a strong component of copepods in fiord communities, krill, when present, are the major pellets producers contributing to the recognizable organic matter reaching the fiord bottom. Surface sediments show negligible seasonal variation in total organic matter, organic carbon and nitrogen, amino acids, and lipids. This may be due to a rapid conversion of sediment material into a pool of sediment microorganisms. The macrobenthos in the deep basin can be dominated by specialized deposit-eaters, such as the echinoderm (Ctenodiscus crispatus), which accumulates fatty acids indicative of an extensive microbial input.

Variability in Time and Space Scaling of Exchange Processes

The pelagic community in fiords is often found to have a higher variability in time and space than that in oceanic water. This is mostly related to differences in advection between the habitats, where the strongest flushing is usually found in fiords. As the physical scale of fiords varies, the balance between internal and external forcing is also likely to vary. The cross-sectional area above the sill is


362

FIORDIC ECOSYSTEMS

therefore an important boundary property, and the ratio between the cross-sectional area and the total fiord volume may indicate the impact of the sill boundary conditions on the fiord. This ratio varies considerably from one fiord to another (Table 2). In a sill fiord, the cross-sectional area of the fiord mouth (A) is believed to constrain the water exchange over the sill. The advective impact has been found to be larger in fiords with a high ratio of cross-sectional area to total fiord volume (see A=V in Table 2). The extent to which the planktonic part of a fiord system is controlled by internal biological processes rather than advective processes depends on the physical scale of the fiord versus the timescale of these processes. As a general simplification we may write eqns [1] and [2]. dB ½1 ¼ rB þ 0:5vRðBB BÞ dt A ½2 V Here B ¼ biomass concentration within the system (mg m3); t ¼ time (s); r ¼ local instantaneous growth rage of B (s1); v ¼ mean absolute current above the sill (m s1); BB ¼ biomass concentration in incoming current (mg m1); A ¼ cross-sectional area above the sill (m2); and V ¼ fiord volume (m3). As B approaches BB , the net advective effect becomes zero. This does not mean, however, that the advective effect has ceased, since the biomass renewal within the system may still be dominated by advection rather than local growth. The growth rate (r) and the advective rate (b ¼ 0:5 vR) have the same dimension (s1) and the ratio r=b41 decides which of the two processes dominates the biomass formation within the system. If r=b41, growth is the dominating process, while r=bo1 indicates advective dominance. The importance of advection relative to the growth of phytoplankton and zooplankton is given R¼

in Figure 3. The much lower growth rate of zooplankton compared with that of phytoplankton implies that the transport influences primarily the zooplankton biomass. Phytoplankton is, however, constrained to the upper, photic, zone where transport processes are most prominent. Zooplankton may utilize the entire water column and the advective influence may thereby be diminished. The zooplankton confined to the advective layer (dotted line) in Figure 3 is influenced three times more strongly by advection than is zooplankton distributed in the entire fiord volume (solid line). Similarly, vertical migration (diel and seasonal) may also reduce the influence of advection for populations depending on the food availability in the advective layer. From eqns [1] and [2] we see that the value of R (ratio between the cross-sectional area above the sill and the fiord volume) gives the order of magnitude of the advective influence in a particular system. This ratio varies considerably from one fiord to another (Table 2), and it may serve as an index indicating the potential advective influence on a system.

Impact of Tidal Currents The tidal amplitude varies by approximately a factor of 2 within the regions where fiords are found. Tidal currents result in no net water exchange when averaged over a tidal period (12.42 h). On the other hand, there is a significant vertical difference in tidal currents, where the highest flow rates are found in surface layer, with concurrent flow in the opposite direction close to the bottom. Therefore, total exchange of organisms is not necessarily proportional to the net exchange of water, and the vertical position of the organisms plays a major role for the advection rate. 5

Table 2 Examples of ratios of cross-sectional area (A) to total fiord volume (V) in Norwegian fiords Fiord

A/V

Advective influence

Linda˚spolls

107

Ryfylke fjord

2 106

Advective influence of Calanus population o internal processes Advective exchange of zooplankton o internal production Advective exchange of zooplankton ¼ internal production Advective exchange of zooplankton biomass 4 internal production Calanus heavily influenced by advective processes

6

Masfjord

8 10

Malangen

4 105

Korsfjorden

104

Advection / production ratio

Zooplankton (0_70 m) 4 3 2

Zooplankton (0_494 m)

1 Phytoplankton (0_70 m) 0 0

5

10 _ Current (cm s 1)

15

20

Figure 3 The relation between the ratio advection/production and current velocity for phytoplankton and zooplankton in a fiord. Modified from Aksnes et al. (1989).


FIORDIC ECOSYSTEMS

Impact of Vertical Behavior The zooplankton community structure in fiords appears to be determined primarily by the species–environment relationship, rather than species–species relationships as in the central oceanic gyres. Different spatial distribution patterns are often observed along the length of the fiord, where populations decrease as they are moved away from their population centers. This along-fiord difference in abundance usually evolves from homogenous low stocks of overwintering populations early the same year. Neritic species prevail at the inner regions, with a downstream decline in the abundance. Most oceanic species have been unable to establish viable populations even in the deepest part of the fiords. However, some, mesopelagic oceanic species have succeeded and the species-specific along-fiord differences in abundance vary and are linked to their differences in vertical behavior.

Most zooplankton migrates on different timescales, and by diel vertical migration (DVM) the population tends to move from deep waters during daytime to surface during night. The residence time during day and night is normally much longer than the transition time between upper and deep waters. Combining the residence time of the animals with the total current in the same depth strata of the water column, a simple prediction of the displacement over time can be estimated. For a species with wide and regular vertical migration, for instance, Chiridius armatus, there is a strong tendency to reside in the same geographical region over time (Figure 4). This species does best at deeper and more oceanic sites, and has been found to reduce its population size by a factor of 4 over a distance of 3–4 km, with an overall decline in abundance toward the head of the fiord. Other species have very low migration amplitude, where ontogenetic migration tends to keep the recruits in surface water for a long time, with an induction of a slow downward migration at older life stages. This is the case for the oceanic species Calanus finmarchicus, which has its population center in the North Atlantic and the Nordic Seas but is a very widely distributed and quantitative species in the fiords of the Northern Hemisphere. C. finmarchicus spends as much as 8–9 months in the deep waters

Deep dwelling

12 8

Migrating between 175 and 225 m

4 0 _4 Displacement (km)

The renewal rate of water and zooplankton can be calculated as a percentage of the total water and zooplankton biomass volume in a fiord with moderate to high tidal amplitudes (Table 3). Comparing contrasting periods in spring and autumn, the total renewal rate of the water is B6% and 14% per day in spring and autumn, respectively. The numbers for the two given size categories of zooplankton are 6% in spring, 3.5% for zooplankton 4500 mm and 12% for zooplankton between 180 and 500 mm in size. The importance of the tidal current for water and zooplankton transport can be estimated by comparing the residual current (i.e., the current velocity after the removal of the tidal component) and the total current. For example, the tidal water is responsible for 23% and 66% of the renewal of zooplankton 4500 mm during spring and autumn, respectively. Equivalent numbers for the smallest size fraction of zooplankton are 19% and 39%. This demonstrates that the tide can play an important role in the transport of zooplankton in fiords.

363

_8 Nonmigrating at 175 m

_ 12

Surface dwelling 12 0 _12 _ 36

Migrating between 10 and 75 m

_ 60 Table 3 Advective daily transport of water and zooplankton biomass in Malangen, expressed as percentage of the entire volume inside the fjord Spring (%)

Autumn (%)

Total

Residual

Total

Residual

6.1 6.6 6.4

3.6 5.1 5.2

14.6 3.5 12.0

8.1 1.2 7.4

_ 84 Nonmigrating at 10 m

_ 108 _ 132 0

Water Zooplankton 4500 mm Zooplankton 180–500 mm

50

100 150 Time (h)

200

250

Figure 4 The travel distance for copepods migrating between 175 and 225 m and residing at 175 m during a 24 h cycle (upper panel), and for copepods migrating between 10 and 75 m and residing at 10 m during a 24 h cycle (lower panel).


364

FIORDIC ECOSYSTEMS

4500 m while reproduction and growth takes place at the surface from April to July. C. finmarchicus is a very prominent species on the shelves and in the fiords during the productive season but descends during summer and early autumn, being drained off from the shallower areas toward overwintering depths in deep basins in the fiords and the continental shelves. In some fiords the biomass build-up can reach 10–20% per day, and the high biomass, in particular in the basins during the autumn, clearly indicates that physical/biological aggregation mechanisms overrule the local production. The mechanisms can be explained by the same simple model given above (Figure 4), where a declining flush rate with depth tends to retain older stages as they initiate the downward migration during the summer. The biomass aggregation will thus continue for several months, draining older stages from the upper and intermediate water layers in the fiord.

Food Web Structure and Functioning The Zooplankton Community

The community structure varies extensively between fiords but reflects mostly the shelf habitats found at similar latitudes (Figure 5). In subarctic waters, the

zooplankton is composed of few species, but with high biomass. Small copepods may be abundant, especially during summer and autumn, and are not major pathways to the juvenile and adult fish. The larger copepods forms (i.e., Calanus finmarchicus and Metridia longa), the chaetognath Sagitta elegans, and the two krill species (Thysanoessa spp.) form easily identifiable trophic links in the transfer of materials to higher trophic levels. They all spawn during spring, matching the spring bloom to variable degrees, and each has a restricted growth period within the time-window from April to October. During the long overwintering period a marked decrease in organic lipid-based reserves takes place in both copepods and krill, accounting for 40–70% of that present at the end of the primary production season. Copepods and krill are often found as soundscattering layers (SSLs) in the basin water of the fiord, and are heavily preyed upon both by demersal and pelagic fish. The Higher Trophic Animals

In fiords there may be around 30 species that have a commercial potential, although only half of these are exploited regularly by man. Some pelagic and demersal fishes are separated into ocean and coastal

er Riv

‘Edible’ deposit feeders demersal fish

Megaplankton _ demersal fish _ ‘refractory’ benthos

Shelf community Diatoms _ large copepods _ planktivorous fish

Fiord community Diatoms _ large copepods _ juvenile fish _ ‘slope’ _ demersal fish

Shallow community Nanoflagellates _ small copepods _ small planktivorous fish or medusae

Figure 5 A generalized structure of the biological community from the shelf to the inner part of the fiord. Modified from Matthews and Heimdal (1980).


FIORDIC ECOSYSTEMS

stocks, where the former spends their time mostly in the ocean and visit the coast during spawning or overwintering. The latter are more confined to the fiords and coastal zones during their entire lifetime. The communities in fiords may have from one to several apex predators. In some fiords cod plays this role, and as many as 70 taxa have been found in cod stomachs from Balsfiorden, northern Norway. Although as many as 15 fish species were identified, krill, capelin, and herring are considered its dominant prey. This emphasizes that cod is weakly linked with the benethic community. Prawns, also, may be more trophically dependent on the pelagial, a feature contrary to the general tendency in the literature to classify them as part of the benthic food chain. The cod stocks in fiords are mostly coastal cod and undertake very little migration. Tagging experiments have documented that individuals are confined to an area from 5 to 8 km in extent during their entire lifetime. Artificially released cod tend to migrate over longer distance and are more easily trapped by fishing gears and other predators, and are thereby subjected to a higher risk of predation. Cod stocks in fiords may have lower growth rate, length, and age at maturity than oceanic stocks, which indicates that they may be subjected to different management strategies in future. Consumption of cod by mammals other than man is substantial, although it varies extensively between fiords. Species such as sea otter, harbor seals, and porpoises do visit fiords for short time-windows, having a more limited affect as predators on the local fish fauna. Cormorants have a less varied diet and are known to roost and feed in fiords from late September to early April, within 8 km of their night roost. Although their toll on local cod stocks may be low, around 1/20 in terms of numbers, compared to cannibalism within the cod population, their local impact is still important since they prefer juvenile cod in the length range 4–50 cm. Other Structural Forces of the Pelagic Community in Fiordic Ecosystems

Light is an important limiting factor for the visual foraging process in fishes, and the light regime may potentially affect the competition between visual and tactile predators. Food demand and risk of mortality are regulated by balance between catching and avoidance between predator and prey, which ultimately may be regulated by visibility in the water column. The seasonal variation in the light may therefore be an important structural force for vertical distributions of important components in the community in fiords. Zooplankton size and density

365

increase with depth; the most visible forms are found only in deep waters. At night, macroplankton and mesopelagic fishes are dispersed in the water column, with a tendency for dispersion of the SSL. All components of the SSL respond to changes in light intensity during day, in order to balance vision versus visibility. The pelagic juvenile fishes in the upper part of the SSL migrate to the surface to maximize their feeding period. The more visible fishes stay in the deeper part of the SSL. At dawn, euphausiids (Meganyctiphanes norvegica) descend from surface waters to midwater depths, and the larger mesopelagic fishes (Benthosema glaciale) and the pelagic prawns (Sergestes arcticus and Pasiphaea multidentata) migrate to even darker water. Large pelagic fishes are found in the entire water column, feeding with the highest densities mainly below and at the deepest end of the SSL. A strong faunal difference is often found between adjacent fiords. This has been linked to the differences in the light climate, fiords with low visibility tending to have a higher component of jellyfish, with the jellyfish being replaced by fish in fiords with higher visibility. This has prompted the hypothesis that the visibility regime may affect the distribution of tactile and visual predators such as jellyfish and fish. The implication is that light is a forcing factor on the marine ecosystem dynamics through the visual feeding process, with potential secondary links to eutrophication. The mechanistic relationship between differences in sea water absorbance and the biological components is not established, but fiords with different visibility regimes will still play an important basis for research needed for ecologically based future management of these important marine biotopes.

See also Beaches, Physical Processes Affecting. Coastal Circulation Models. Copepods. Fish Larvae. Fish Migration, Vertical. Fish Predation and Mortality. Krill. Macrobenthos. Marine Mammal Trophic Levels and Interactions. Meiobenthos. Mesopelagic Fishes. Patch Dynamics. Pelagic Fishes. Phytoplankton Blooms. Primary Production Processes.

Further Reading Aksnes DL, Aure J, Kaartved S, Magnesen T, and Richard J (1989) Significance of advection for the carrying capacities of fiord populations. Marine Ecology Progress Series 50: 263--274.


366

FIORDIC ECOSYSTEMS

Falkenhaug T (1997) Studies of Spatio-temporal Variations in a Zooplankton Community. Interactions between Vertical Behaviour and Physical Processes. Dr Scient thesis, University of Troms; ISBN 82-91086-12-5. Howell B, Moksness E, and Svsand T (eds.) (1999) Stock Enhancement and Sea Ranching. Oxford: Fishing News Books.

Matthews JBL and Heimdal BR (1980) Pelagic productivity and food chains in fiord systems. In: Freeland HJ, Farmer DM, and Levings CD (eds.) Fiord Oceanography. New York: Plenum.


FISH ECOPHYSIOLOGY J. Davenport, University College Cork, Cork, Ireland Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 910–916, & 2001, Elsevier Ltd.

Introduction The earliest known jawless fish (Agnathans) date from the Cambrian period 500–600 million years ago. It is generally agreed among paleontologists that they evolved from sessile, filter-feeding ancestors in shallow fresh or brackish water, not the sea. Their environment was characterized by low salt levels and was probably turbid and at least intermittently oxygen-depleted (hypoxic). They have since spread to virtually all aquatic habitats (Table 1), but freshwater is still a stronghold, with a third of all fish species living in rivers, lakes and streams even though such habitats only contain a minute proportion (o0.01%) of the total water on earth. Fish make up the most numerous vertebrate class, both in terms of species’numbers and biomass. They form a remarkably plastic group, exhibiting a diversity of sizes (from 8 mm Philippine gobies, Mistichythys

luzonensis, to 18 m whale sharks Rhincodon typicus), shapes and life histories. Living forms include the numerically dominant teleost bony fish, the elasmobranchs (sharks and rays) and smaller groups such as lungfish, coelocanths (Latimeria spp.) and the surviving jawless lampreys and hagfish. Despite this great diversity, fish are monophyletic, i.e. they have a single origin. All living fish share anatomical and physiological features of their earliest ancestors and are still constrained by them to a greater or lesser extent. Basic fish features include the following:

•

• • • •

Table 1 Relative proportions of c. 30 000 living fish species living in different habitats Major category

Subcategory

% of total

Marine fish (58.2%)

Shallow warm water Shallow cold water Deep benthic Deep pelagic Epipelagic Diadromous Primary inhabitants Secondary inhabitants

39.9 5.6 6.4 5.0 1.3 0.6 33.1 8.1

Freshwater fish (41.2%)

*

*

*

*

A third of species are primary inhabitants of freshwater, though freshwater makes up only 0.0093% of the total water of the earth. Possibilities of isolation and allopatric speciation are greater in freshwater than in seawater. Many marine species have re-invaded freshwater, or are at home in both media. Several intertidal and freshwater fish species spend some of their time on land.

After Cohen DM (1970) How many recent fishes are there? Proceedings of the California Academy of Science 38: 341–345.

Vertebral column with associated myotomal (segmented) muscles. The vertebral column sets fish length and prevents shortening of the animal when muscle contraction powers undulatory swimming. Head with food capture apparatus and bilaterally symmetrical sense organs. Gills for respiratory exchange, excretion of nitrogenous waste, plus regulation of body fluid ion content and pH. Closed vascular system with chambered heart, circulating red blood cells and (in teleost fish) dilute blood plasma (250–600 mosmol kg1, compared with 1000 mosmol kg1 of seawater). Elasmobranch fish and coelocanths have similar plasma ionic levels to teleost fish (i.e. much lower than the ionic concentration of seawater), but have high urea and trimethylamine oxide (TMAO) levels that result in total plasma osmolarity being similar to that of seawater (Figure 1).

Biotic and Abiotic Factors in Distribution of Marine Fish Distribution in fish is always controlled by a mixture of biotic and abiotic factors. For example, herbivorous marine fish are limited to shallow water where photosynthesis by microalgae, seagrasses or seaweeds is possible. Carnivorous fish (particularly postlarval and juvenile forms) may also be limited to such areas because they specialize in feeding on herbivorous invertebrates. Young gadoid fish are found living for protection beneath the bells of stinging jellyfish – a resource limited to near-surface waters at specific times of the year. These are all biotic constraints. Abiotic influences are those imposed by physical or chemical factors such as temperature, salinity or oxygen tension, and they are the main


367

368

FISH ECOPHYSIOLOGY

_1

Plasma/hemolymph osmolarity (mosmol kg )

1000

Urea / TMAO

500 Ions

Elasmobranchs

Teleost fish

Hermit crabs

Seawater

0

Figure 1 Diagrammatic comparison of relative contributions to total plasma/hemolymph osmolarity of ions and urea/ trimethylamine oxide (TMAO) in teleost fish, elasmobranchs and hermit crabs (Pagurus bernhardus). Data for seawater given as reference.

focus of this article since they profoundly affect fish physiology and tolerances and therefore distribution. There are other features that are strictly abiotic, in particular physical habitat type (e.g. mud, sand, coral reefs, tangled mangrove roots), but they do not impinge directly on fish physiology, so will not be considered here.

Tolerances and Limits to Distribution Presence of Water

Limitation to an aqueous habitat is the most fundamental physiological constraint imposed on fish. No elasmobranch or agnathan species can survive out of water, and only a few dozen amphibious teleost species plus the three surviving lungfish species have the ability to live out of water for significant periods. Most of these species are freshwater, so outside the scope of this encyclopedia. Amphibious intertidal fish such as mudskippers (Periophthalmus sp.), shannies (Lipophrys pholis) or butterfish (Pholis gunnellus) are rarely more than a few meters from

seawater and are emersed only for a few hours at a time. Eels (Anguilla sp.) are catadromous and move between freshwater and seawater, these migrations often involving movement on damp/wet land between water bodies. When fish are emersed they have problems of respiration, excretion, acid–base balance and locomotion. Fish respiration involves the passage of an incompressible medium (water) over the gills. If an unadapted fish is taken out of water it becomes short of oxygen (hypoxic), accumulates CO2 (hypercapnia) and cannot excrete Hþ – all because the gills collapse and the animal cannot circulate compressible air over them. It soon dies because the blood becomes acid, not because of lack of oxygen, and usually death will occur long before dehydration becomes a factor. Mudskippers and shannies have strengthened gills with relatively few filaments that are less prone to collapse. They also have scaleless, well-vascularized skins that allow respiratory exchange; the gills are relatively less important for respiration, though they can still take up oxygen from water held within the buccal cavity, and are involved in Hþ regulation. Nitrogen excretion is also a problem for amphibious fish. Most fish excrete nitrogen as NH3 or NH4 þ ; with 50–70% of excretion being by diffusion across the gills in marine fish. Although NH3 =NH4 þ is metabolically inexpensive to produce, it is toxic and therefore cannot be accumulated within the body. Amphibious marine fish such as mudskippers reduce protein breakdown when in air, and also accumulate a proportion of N2 as urea or trimethylamine oxide (TMAO), both of which are less toxic than ammonia. However, for marine fish, nitrogenous excretion is a fundamental constraint on survival on land; they have to return to the sea to get rid of accumulated nitrogen and metabolites. Basic fish anatomy is unsuitable for terrestrial locomotion, though butterfish and eels ‘swim’ through three-dimensional habitats, such as pebbles and thick grass, relying on secretion of mucus for lubrication. Shannies and mudskippers have strengthened prop-like pectoral fins that stop them falling over and raise the belly off the ground to some extent. They are propelled in a series of hops by the tail. Depth of Water

Four decades ago Jacques Piccard took photographs from the bathyscaphe Trieste of tripod fish (Chlorophthalmidae), resting on the bed of an oceanic trench at a depth of over 10 000 m. In doing so he demonstrated that fish could live at all depths, despite their shallow-water origin. Trawling, cameras


FISH ECOPHYSIOLOGY

and submersible observations have confirmed that a diverse ichthyofauna may be found at all depths, in all seas. However, increasing depth may control the sort of fish that are found. Some depth-related constraints are biotic; deep water generally has a restricted energy supply due to absence of light and distance from the productive surface layers. However, there are two major physical problems imposed by depth: increasing pressure and low temperature (particularly at depths greater than 1000 m). In addition, there is an important related chemical problem. In warm surface waters, the sea is practically a saturated solution of calcium carbonate and relatively little energy is needed to maintain calcareous materials (e.g. bone, shells) in solid form. However, solubility rises with increasing pressure and decreasing temperature. In consequence, building and sustaining solid calcareous materials becomes more expensive, particularly at depths beyond 3000 m (below which calcareous sediments are unknown). The problems of pressure and calcium carbonate solubility interact in the physiology of fish buoyancy. A high proportion of shallow-water teleost fish have swim bladders, which develop from gut diverticula. They have sophisticated volume regulatory mechanisms (employing the lactate-secreting gas gland and its associated countercurrent multiplier system), particularly in physoclistous fish in which the swim bladder is isolated from the gut. The original adaptive value of swim bladders probably lay in offsetting the burden of dense scales and armoured skulls, so that shallow-water benthic fish could swim into the water column without undue effort. At depths down to about 1000 m some 75% of fish have swim bladders, but at greater depths pelagic fish usually have no swim bladder or a swim bladder filled with fat; they also show progressively reduced musculature and ossification, plus a very high water content. Lack of ossification has been interpreted as an energy-saving strategy in a foodpoor environment where dissolution of calcium carbonate takes place. Loss of swim bladders has been attributed to the metabolic expense at great depth of pumping gas into swim bladders. Interestingly, benthopelagic fish (i.e. those living close to the sea bed) from deep water can possess working swim bladders even at depths of 7000 m; they are also of robust skeleton and musculature. The near-sea bed environment is now known to be much more energyrich than the water column above, particularly due to the fall of large carrion items. This suggests that benthopelagic fish can ‘afford’ to expend energy counteracting abiotic depth-related factors on fish form, because of the biotic influence of a good food supply.

369

Temperature

Fish are almost all ectothermic animals with no significant production or retention of metabolic heat. A few tuna species can keep their locomotory muscles warm, and some big lamnid sharks (e.g. the great white shark, Carcharodon carcharias) are partially endothermic with core body temperatures being held at around 251C in waters of 151C. However, the body temperature of most fish is directly determined by environmental temperature. Metabolic rate in ectotherms is strongly affected by temperature, with a useful rule of thumb (the Q10 relationship) stating that metabolic rate is doubled by an increase of 101C in environmental temperature. There is a huge literature devoted to the effects of temperature on aspects of the physiology, development and ecology of marine fish. However, despite this wealth of information, the question of whether thermal physiological constraints control fish distribution is difficult to answer, since marine fish are found in all available habitats, from Antarctic ice tunnels at 2.51C to Saudi reef pools at over 501C. Antarctic fish usually die at around þ 51C, whereas most temperate fish are stressed severely at temperatures above 301C. Tropical fish reach their thermal limits at about 451C (also the upper thermal limit for most tropical marine invertebrates). The position and breadth of a species’ thermal niche is determined by a variety of factors. At the biochemical level, fish must have enzymes with appropriate thermal optima. Cell membranes are effectively liquid crystals that must remain fluid if the cell is to survive. Maintaining fluidity over the full environmental temperature range requires modulation of the fatty acid composition of membrane lipids (homeoviscous adaptation). In many fish these characteristics can vary geographically and seasonally, but within overall limits which are species specific. These limits do constrain fish distribution. In the northern hemisphere capelin (Mallotus villosus) do not penetrate much further south than the 51C summer surface isotherm, whereas the corresponding limit for cod (Gadus morhua) is about 201C. In marked contrast, deep-water fish living at very stable low temperatures (c. 21C) can have extremely wide distributions despite limited thermal tolerance. For example, the orange roughy (Hoplostethus atlanticus) living at 800–1200 m has been fished off New Zealand, Namibia and northern Europe. A specialized warm-water example of thermal constraints on distribution is provided by flying fish (Exocoetidae). Flying fish take off through or from the sea surface at very high speed (15–30 body lengths s1). This in turn demands extremely high


370

FISH ECOPHYSIOLOGY

rates of tail beat (c. 50 beats s1). Calculations demonstrate that take-off is unlikely to be possible at water temperatures below 201C, and surface water of this temperature approximates to the northern and southern limits of this essentially tropical group. Physiologically, fish need greater gill areas per unit mass at high temperature because their respiratory requirements are greater and because the solubility of oxygen decreases with rising temperature, so less environmental oxygen is available. At the ecological level, high temperatures imply high metabolic rates and elevated food consumption. A typical tropical fish needs roughly six times the oxygen to support resting metabolism as does a typical polar fish. High temperature, high activity lifestyles can only be sustained in energy-rich environments. However, it should not automatically be assumed that low-temperature environments (e.g. the deep sea, polar waters) are inexpensive to live in. Low temperature is associated with high viscocity so swimming demands proportionally more energy, as does pumping of blood around the body. Antarctic fish in particular tend to show viscocity-related modifications; they (and some deep-water bathypelagic fish) generally have low hematocrits (i.e. have relatively few red blood corpuscles) so that effective blood viscocity is reduced. Icefish (Channichthyidae) provide extreme examples of this, having no red blood corpuscles or hemoglobin and possessing wide-bore blood vessels through which viscous plasma flows more easily. The trade-off is that they are sluggish fish with extremely low endurance, demonstrating again that lifestyle can be constrained by temperature. Oceanic seawater of salinity 34% (osmolarity 1000 mosmol kg1) freezes at about 1.91C. Elasmobranch fish are at little risk of freezing because their blood has a similar osmolarity (Figure 1). Because teleost fish have relatively dilute body fluids (300–600 mosmol kg1) they are potentially liable to freezing at temperatures of 0.6 to 1.01C, when seawater is still fluid. Lower temperatures than this occur in the winter in the Arctic and throughout the year in the Antarctic. This contrasts with the situation for freshwater fish in which freezing cannot take place unless the water around them is itself frozen. Intertidal pools can become even colder, unfrozen high salinity water beneath ice sheets reaching 3 to 81C. High latitude marine fish show various adaptations to deal with freezing risk. Anadromous arctic charr (Salvelinus alpinus) migrate from the sea to freshwater in the winter, whereas other arctic fish migrate to low latitudes or deep water. There appears to be a fundamental constraint on the distribution of resident intertidal fish, which are absent from northern Norway, Russia and Canada

throughout the year – presumably because they cannot compete effectively in deeper water in winter, are too small to migrate significant distances southwards, and cannot avoid freezing. Antarctic fish and permanent surface-water residents of the winter Arctic all exhibit physiological adaptations that permit freezing avoidance even when pack ice is present. All have relatively high blood osmolarities, depressing their potential freezing points, and most can produce so-called antifreezes: peptides or glycopeptides. These molecules do not actually stop ice formation in body fluids, instead they adsorb onto the crystal surfaces of minute ice nuclei and prevent these nuclei from growing or propagating. In icefish these mechanisms are effective to 2.51C, allowing them to live in ice tunnels in Antarctic ice shelves. Several arctic species overwinter in deep water which is colder than their potential blood freezing point, but does not contain ice nuclei that can initiate freezing. This supercooled state is precarious and there are records of schools of capelin (for instance) straying into shallow water containing ice during cold weather and freezing instantly. Salinity

The great majority of marine fish species live under very stable salinity conditions (34–35%; osmolarity c. 1000 mosmol kg1). This medium, though stable, is much more concentrated than the freshwater or brackish media encountered by ancestral fish. As a result, the osmotic and ionic physiology of marine teleost fish is very different from that of freshwater fish, and relatively few species can live in both media (mainly anadromous and catadromous fish such as salmonids and eels). Elasmobranch fish (sharks, skates and rays) are adapted to seawater by virtue of high blood urea levels that make the blood slightly hyperosmotic to seawater so that they have little or no osmotic problem. Although this group does have a few brackish water species, it is generally limited to fully marine habitats and will not be discussed further. Briefly, marine teleost fish have to drink seawater to replace water lost osmotically (mainly across the gills) because the blood is much more dilute than seawater (Figure 2). To gain access to the water taken into the foregut, salt pumps (actually ATP-ase enzymes embedded in cell membranes) sited on the intestinal wall pump salts from gut fluid to blood until (in the posterior parts of the gut) the osmolarity of the gut fluid is below that of the blood, so that water flows into the blood osmotically. The combination of drinking seawater and active desalting of the gut fluid results in much salt uptake, augmented by diffusion from the


FISH ECOPHYSIOLOGY

Medium 1000 mosmol

371

H2O

H2O

H2O

Na+ Na+ Na+

Drinking

Intestine Na+ Na+ Na+ Na+ Na+

Blood 300 mosmol H2O

800 mosmol 400 mosmol 250 mosmol

1000 mosmol

Na+ Na+ Na+ Na+ Na+ Na+

Gills

H2O

Na+ Stomach +

Na H2O

H2O

H 2O

Active salt pumping Osmotic loss of water Osmotic uptake of water

Figure 2 Diagrammatic representation of osmotic and ionic regulation in a marine teleost fish.

salty external medium. To counteract this, salt pumps located mainly in ‘chloride cells’ on the gills actively pump salt outwards. In the bulk of the world’s oceans, salinity does not constrain distribution. Difficulties only occur at the seas’ margins, in estuaries, lagoons and pools where salinities can be much higher or lower. Generally, there appears to be an upper limit of survivable salinity for fish of around 80–90%, with one or two specialists (such as the killifish Fundulus heteroclitus) tolerating up to 128%. Impressive though this performance is, it is much poorer than that exhibited by many invertebrates, particularly crustaceans, some of which can tolerate up to 300% (e.g. the brine shrimp, Artemia). Fish are constrained by what is known as the ‘osmorespiratory compromise’. Fish gills are necessarily large in surface area to support gaseous exchange. Gill epithelia are also thin to permit ready diffusion. Unfortunately, these two characteristics also favor rapid osmotic loss of water that has to be replaced by drinking, plus salt diffusion that must be opposed by salt pumps. There comes a point where the balance breaks down; killifish are unusual in that they tolerate a 30% rise in plasma osmolarity at 128%, most fish would succumb. The majority of marine fish have blood osmolarities of around 300 mosmol kg1, equivalent to about 10%. If unadapted marine fish are placed in media less concentrated than this, they die because of blood dilution caused by osmotic uptake of water and diffusional loss of salts. To survive in dilute media, euryhaline fish of marine origin have to stop drinking the medium and pump salts inwards at the gills, effectively reversing the osmotic physiology exhibited in seawater. Many fish of this type acclimate slowly over a period of days to dilute media, since profound microanatomical and biochemical changes have to

take place; migratory salmonids fall into this category. Some fish that forage regularly into brackish water (flounders (Pleuronectes flesus), mullets (Mugilidae), ro`volos (Eleginops sp.)) respond more quickly, and a few species such as the shanny are capable of reacting to extreme salinity changes in a matter of minutes. Shannies inhabit the intertidal zone and may be found in crevices that are fed by freshwater runoff when the tide falls. They exhibit the constrained features of such highly euryhaline fish, i.e. small size, thickened gills and a low skin permeability to salts and water, that slow changes in blood concentration and reduce the energetic costs of regulation. Most marine fish that are regularly exposed to low salinity are benthic and slow moving, another consequence of the osmorespiratory compromise. Oxygen Tension

Fish in general, and marine fish in particular are intolerant of oxygen-depleted (hypoxic) conditions. Many freshwater habitats are hypoxic and fish have multiple hemoglobins to deal with this situation and extract oxygen from hypoxic water. Broadly speaking, marine habitats are mostly close to equilibration with the atmosphere (air-saturated; normoxic), so oxygen tension poses no constraints and most fish have few or single hemoglobins. Pollution incidents have often revealed the sensitivity of marine fish to hypoxia, with diatom bloom formation sometimes resulting in massive fish kills at night when plant respiration has dramatically reduced water column oxygen tension (Po2). In general marine fish avoid hypoxic areas rather than tolerating them, though large schools of clupeoids (e.g. herring, Clupea harengus) create their own hypoxic environments in the heart of the shoals. Shannies will even leave nocturnally hypoxic


FISH ECOPHYSIOLOGY

intertidal pools and respire in air, particularly in summer when their respiratory demands are high. Marine fish unable to avoid transient hypoxic conditions (e.g. sole, Solea solea in shallow organic-rich water) usually respond by reduced activity, depressed basal metabolic rate and activation of anaerobic metabolism. This is a short-term response and is supplemented by behavioral responses such as burstswimming to the surface where higher oxygen tensions prevail.

100%

Performance (e.g. growth)

372

Lower lethal limit

Upper lethal limit Optimum

50%

Tolerance 0% Low

High Temperature

pH

Fish have a plasma pH of 7.4–8.1. Unlike higher vertebrates they have limited internal buffering mechanisms and the viscocity of water means that ventilatory pH control is much less effective than in air-breathers. Excretion rather than respiration controls fish acid–base balance. In many freshwater habitats, whether natural or affected by acid rain, environmental pH poses considerable physiological costs and constraints. This is not the case for marine fish because the sea is a slightly alkaline environment (pH 7.8–8.0) that has enormous buffering capacity for Hþ and CO2 and poses no problems whatsoever. Only in rock pools are great pH fluctuations known (7.2 at night; 10.6 by day), and even here there are no documented problems for rock pool fish.

Optima Optimal abiotic environmental conditions for fish species have been studied from two perspectives. First, there are now numerous commercially valuable marine fish species in culture, principally for human food, but increasingly in tropical countries for the aquarist trade. An extensive literature reporting on the ideal conditions for survival, rapid growth and effective reproduction has arisen. Much of the work done has involved multifactorial experimental approaches (e.g. combinations of temperature, salinity and oxygen tension) and these have often revealed changes in optima at different stages of the life history. Rearing densities are high and have biotic effects on optimal rearing conditions, but these in turn can create constraining abiotic problems (e.g. NH4 and nitrite accumulation) that do not arise in nature. Such studies have revealed constraints on where economic acquaculture can be performed. For example, halibut (Hippoglossus hippoglossus) farming is practiced only in cold areas of northern Europe (e.g. Norway, Scotland) because Hippoglossus is a deep-water species intolerant of elevated temperature. Conversely, turbot (Scophthalmus maximus) farming, once tried in such areas, proved to be

Figure 3 Diagram to illustrate tolerance range, optimum performance and lethal limits (using temperature as an example).

uneconomic because extra (expensive) heat was needed to secure fast growth; production is now centered in France and Spain. Secondly, abiotic optima have been considered from a biogeographical perspective, almost exclusively in thermal terms (Figure 3). This is linked with the constraints on distribution already discussed, but is presently a rather poorly developed research area, ripe for development. For temperature, there is evidence that fish seek out preferred (assumed optimal) temperature regimes, though much of this work has been conducted on freshwater species. Widely distributed marine species have been assumed to have optima either in the center of distribution, or close to the warmer limits, but some work on fish living at the edge of distributions show no evidence of maladaptation or of particular population instability. Difficulties lie in deciding what optimal performance means, and in disentangling biotic and abiotic influences. Fish in temperate zones usually grow more quickly in the warmer areas of their distribution. They may reproduce at an earlier age, often at smaller size. How much of this is due to biotic influences (e.g. quantity and quality of food supply, trophic structure of the local ecosystem) is usually unclear. Optimal performance essentially involves maximizing contribution to the gene pool; establishing the ecophysiological conditions that deliver this is still a challenge.

See also Antarctic Fishes. Deep-Sea Fishes. Eels. Fish Larvae. Intertidal Fishes. Salmonids.

Further Reading Bone Q, Marshall NB, and Blaxter JHS (1999) Biology of Fishes 2nd edn. Glasgow: Stanley Thomas.


FISH ECOPHYSIOLOGY

Dalla Via D, Van Den Thillart G, Cattani O, and ortesi P (1998) Behavioural responses and biochemical correlates in Solea solea to gradual hypoxic exposure. Canadian Journal of Zoology 76: 2108--2113. Davenport J and Sayer MDJ (1993) Physiological determinants of distribution in fish. Journal of Fish Biology 43(supplement A): 121--145. DeVries AL (1982) Antifreeze agents in coldwater fish. Comparative Biochemistry and Physiology 73A: 627--640. Kinne O (1970) Marine Ecology, vol. 1. New York: Wiley Interscience.

373

Kinne O (1971) Marine Ecology, vol. 2. New York: Wiley Interscience. Marshall NB (1979) Developments in Deep Sea Biology. Poole: Blandford Press. Rankin JC and Davenport J (1981) Animal Osmoregulation. Glasgow: Blackie. Sayer MDJ and Davenport J (1991) Amphibious fish: why do they leave the water? Reviews in Fish Biology and Fisheries 1: 159--181.


FISH FEEDING AND FORAGING P. J. B. Hart, University of Leicester, Leicester, UK & 2009 Elsevier Ltd. All rights reserved.

Introduction There are approximately 24 600 species of fish of which 58% or 14 268 live in the sea. The sea covers c. 71% of the surface area of the Earth and has an average depth of around 3800 m, so that the total volume of the marine environment is about 1370 106 km3. Much of this volume, removed from the influence of the sun’s rays, is an inhospitable place to live, being dark, cold, and very low in available food. With such a volume for living, it is no surprise to learn that there are some 15 basic ways in which fish can gain food from the environment (Table 1). Because the open ocean and the deep ocean have low productivity compared with the shallow seas, most of the fish diversity is found in waters less than 200-m depth with the highest concentrations being found in tropical waters over coral reefs. The fish in these areas also have the greatest diversity of ways of making a living. Coral reefs and other inshore areas also have the most complex food chains with many links between fish and their prey.

(Melanogrammus aeglefinus). The related rat-tail macrourid living at 3000-m depth is more likely to use olfaction or the lateral line to find prey, and many deep-sea fish have very large mouths to allow them to take whatever prey they encounter. The categories in Table 1 are a useful way to illustrate how form, function, and behavioral habits influence the characteristics of different feeding types. Fish biologists also use the term ‘guild’ when classifying feeding modes of fish. In contrast to a niche, a guild refers to a collection of fish species, which can be from different taxonomic groups that feed on the same type of prey and show convergent adaptations to the food. For example, herring and mackerel are pelagic planktivores and show many similar adaptations of body form that are designed to cope with living in the open sea, yet the two species belong to very different taxonomic groups. Table 1 fishes

Feeding modes of fishes: Major trophic categories in

Category

Examples

Detritivore Scavengers

Mullet, Mugil Dogfish, Squalus; hagfishes, Myxinidae

Herbivores Grazers Browsers

Modes of Feeding in Fishes During the course of evolution, fish in the marine environment have developed a diverse array of behavioral, morphological, and physiological adaptations to cope with the food they most commonly eat. Although fish feeding habits can be classified into a relatively few groups, the diversity within each group is significant. The different modes of feeding are shown in Table 1 together with a selection of illustrative species. With this table in mind, it becomes possible to examine in more detail the principles of behavioral adaptations used to cope with different conditions. Feeding mode can be classified by the type of food eaten. A species adopting a particular type of food, say that of a piscivore, will develop a body form and a set of foraging tactics suiting it to the particular types of prey taken and the habitat in which the piscivore lives. For an example, a whiting (Merlangius merlangus) living in shallow areas of the North Sea uses vision to locate prey and has a larger mouth than an invertebrate feeder such as the haddock

374

Phytoplanktivores Carnivores Benthivores Picking at relatively small prey Disturbing then picking at prey Picking up substratum and sorting prey Grasping relatively large prey Zooplanktivores Filter feeders Particulate feeders Piscivores Ambush hunters Lurers Stalkers Chasers


Parrotfishes, Scaridae Surgeon fishes, Acanthuridae Peruvian anchoveta, Engraulis ringens

Lemon sole, Microstomus kitt Gurnards, Triglidae Black surfperch, Embiotica jacksoni Triggerfish, Balistes fuscus

Menhaden, Brevoortia tynrannus Anchovy, Engraulis mordax Megrim, Lepidorhombus wiff-iagonis Angler fish, Lophius piscatorius Trumpet fish, Aulostomus maculatus Bluefin tuna, Thunnus thunnus

375

FISH FEEDING AND FORAGING

Most shallow-water environments contain varying amounts of detritus derived from dead plants and animals. This can provide a source of food for some fish, although this mode does not comprise a major feeding type in the sea when compared with fresh waters. In Table 1, mullets (Mugilidae) are given as an example and this highlights a problem with the classification: the categories are not exclusive. Very few fish specialize to such a degree that they never eat anything but the prey type classed as their principal food. So, even though mullet species do eat detritus, they also graze on plant material and capture invertebrate food. It is probably true of all species that eat some detritus that this food source is a supplement to their diet, resorted to when other items are scarce. A large range of fish types feed on the dead remains of other fish, marine mammals, or invertebrate species. Hagfish (Myxinidae) are primitive and have no jaws. Inside their buccal cavity they have teeth that are used to rasp flesh once they have attached themselves with the suckerlike mouth. Although they often feed on dead fish, they also consume living fish if they can first obtain a good hold on them. This is facilitated by the presence of an irregularity on the skin of the prey, such as a wound. Other species, such as the spur dog, Squalus acanthias, take dead material if it is available, although they mainly eat fish and larger invertebrates. As with many carnivores from other animal groups, dead meat is rarely ignored. Herbivores in the sea are limited in their choice. They can either frequent the shallow waters and consume macroalgae or algae encrusting rocks, or they can live in the open water near the surface and eat phytoplankton. If they choose the second option, they are most likely to eat zooplankton also. Grazers and browsers are most common on coral reefs (Table 2), where there are numerous species feeding on algae. Many herbivores are very selective in the species they eat and the grazing effect has a strong influence on competition for space between algal species. Herbivorous fish on coral reefs adopt one of three feeding strategies: they defend a territory, with some species of damsel fish (Pomacentridae) ‘tending’ gardens; they can adopt a home range within which all feeding occurs, as exemplified by some species of pomacanthid angelfishes; or they feed in mixed species groups as in some surgeon fishes. Herbivores on a reef feed only during the day and hide in crevices during the night. Although they are separated in Table 1, phytoplanktivores and zooplanktivores will be dealt with together. Fish that feed on plankton can adopt one of two tactics: either they can sieve the water to extract

the plankton or they can pick off items individually. The two are presented as individual tactics in Table 1, but in reality species will switch between the two depending on the density and size of food. For species focusing on phytoplankton, sieving is the only alternative, as the plants are too small to take individually. The Peruvian anchoveta (Engraulis ringens) takes a mixture of phytoplankton and zooplankton but, because phytoplankton is so rich, the bulk of what they eat could be of plant origin. Most other planktivores eat a mixture of both, with zooplankton predominating. The classic planktivores are species such as the herring (Clupea harengus), mackerel (Scomber scombrus), pilchard (Sardina pilchardus), and sprat (Sprattus sprattus). They live in the epipelagic region of the ocean, have fusiform streamlined bodies, and most often live in large shoals. Many of them make significant migrations to reach feeding areas that are seasonally worth exploiting. An example is the mackerel stock that spawns off southwest England in the spring and then migrates into the North Sea either via the west coast of Britain or up through the English Channel. A wonderful example of a plankton feeder is the basking shark (Cetorhinus maximus). It is remarkable that such large animals, up to 10 m long, can be sustained by their microscopic prey. To survive, these 3000-kg fish have to filter very large volumes of water and do so by swimming for long periods with mouths wide open. Fine rakers on the gill arches act

Table 2 Proportions of different types of feeders in a temperate and a tropical marine system Feeding category

Herbivores Phytoplankton Benthic diatoms Filamentous algae Vascular plants and seaweeds Detritivores Carnivores Zooplanktivores Benthic invertebrates Piscivores Omnivores a

Category absent.


Gulf of Maine, Atlantic

Marshall Island, Pacific (coral reef)

%

%

N

N

0.7 0 0

1 0 0

0 1.5 16.0

0 3 33

0

0

8.7

18

0.7

1

3.9

8

16.9 41.2

25 61

6.3 54.9

13 113

39.2

58

a

2.0

3

8.9

18

376


as filters removing plankton from the stream of water leaving the gill slits. In planktivores such as the herring, prey items are mostly selected and the frequency of prey species found in the stomach is not the same as their frequency in the environment. A famous study of the diet of herring by Sir Alister Hardy, made in the early 1920s, showed how complex the feeding habits of a fish are. Like all species of teleost fish, herring grow throughout their lives, starting as microscopic larvae and finally reaching a size of around 30–40 cm. As revealed by Hardy, the diet of the fish changes dramatically as the fish increases in size, and the figure that Hardy produced to show this (Figure 1) has become a classic of the marine biology literature. As larvae, the herring feed on very small planktonic prey such as the early stages of copepods, larval mollusks, tintinnids, and dinoflagellates. At this stage of their lives, herring are as much food for other fish as they are predators themselves. With growth, the young herring can begin to take larger planktonic prey such as the copepods Pseudocalanus, Temora, and Acartia, common in inshore waters off the British Isles. Juvenile and adult herring feed extensively on Calanus finmarchicus, one of the most common copepods, euphausiids (krill), amphipods, and fish. By

Young herring

7_12 mm

12_ 42 mm

Tomopteris

Medusae

Sagitta

changing their diet through their life history, the herring are moving niche too, and this also has a spatial component as the young herring live in nursery areas close inshore. Carnivorous fish come in a wide range of forms (Table 1). A basic division is between species that feed mainly on prey dwelling in or on the bottom and those that take prey from the water column. Benthic feeding fish show a range of adaptations reflecting the differences in lifestyle of the species. For example, the lemon sole (Microstomus kitt) is a visual feeder swimming over the bottom searching for annelid worms, which make up the bulk of its diet. In contrast, the sole (Solea solea) lies buried in the sand or mud during the day and forages only at night or when light conditions are very low during the day. It searches for food by touch. Both species could be categorized as ‘benthivores’ that pick at relatively small prey, but their differences are not insignificant. These differences are largely behavioral as both species are bottom dwellers superbly designed for their habitat, having flattened bodies and the habit of burying themselves in sediment. As with herring, bottom feeders show life history changes in feeding behavior. For example, the black surf perch (Embiotica jacksoni), living on reefs off

42_130 mm

Adult herring

Limacina

Oikopleura

Pleurobrachi Ammodytesjuv Cypris balanus larvae

Larval mollusca

Pseudo calanus

Decapod larvae

Hyperiid amphipods

Evadne Acartia

Temora

Podon

Calanu

Tintinnopsis

Peridinium

Diatoms and flagellates (plants)

Figure 1 The food of herring from larval stages to adults. Also shown are the connections between the prey of herring and the food they eat. Redrawn from Hardy AC (1924) The herring in relation to its animate environment. Part I: The food and feeding habits of the herring. Fishery Investigations, London, Series II 7(3): 53.



southern California, is named as an example in Table 1 of a species that picks up substratum and, in the mouth, sorts prey from gravel and sand. These fish use this ‘winnowing’ tactic only once they have grown above a certain size. When they are small they fall into category ‘Picking at relatively small prey’ (under ‘Benthivores’) of Table 1 as they feed by picking up each prey item. Their diet is limited, by the gape of the mouth, to food particles below a certain size. The greatest problem for a piscivore is that its prey is mostly mobile and well able to see the predator coming. Two basic tactics are used by piscivores to capture prey: either they use stealth in various forms or they try to outswim the prey in a chase. These tactics have had profound influences on the selective forces influencing the fish found in each group. Those that use stealth have developed along two routes. Many species have special adaptations to lure prey close; the classic example of this is the angler fishes (Lophiidae) with their dorsal fin ray modified to form a movable rod with a lure on the end. These fish have also developed body coloration and shapes that camouflage them as they sit and wait on the bottom. Ambush hunters have used other means of getting close to prey. They hide in weed or adopt coloration that makes them inconspicuous. For example, the megrim (Lepidorhombus wiffiagonis) eats mainly fish and shows the characteristic morphology of a piscivore despite its flattened body form. It has a slim body relative to other flatfish and has large eyes and mouth. To catch fish it lies half-buried on the bottom until a fish comes close, when it springs forward to make a capture. Stalkers are also well camouflaged but get close to their prey by either using cover to disguise their intentions or by moving so slowly that the prey do not notice the advance until too late. For example, the trumpet fish (Aulostomus maculatus) living over coral reefs join shoals of nonpredatory fish as a way of coming close to their prey. The remarkable aspect of this tactic is that the trumpet fish changes its head color to match the color of the fish in the shoal. At sea, fish such as tuna (Thunnus spp.), the larger sailfish (Istiophorous albicans), and salmon (Salmo salar) hunt their prey at speed. These fish have bodies designed for fast swimming and also have large mouths, often with backward-pointing teeth, to grab the prey securely once caught. Many prey fish that are attacked in this way have developed behavioral tactics to reduce the risk of predation. They can shoal or school, they can develop camouflage, or they can avoid contact with predators by appearing only at night. A few species of fish have specialized in exploiting others for food. At its most aggressive, this mode

377

includes fish that eat scales or fins of other species, although most of the examples of these modes are from fresh water. Some species have adopted the role of cleaners who specialize in picking ectoparasites off other fish. This mode is not exploitative in that both sides of the interaction benefit. There are wrasse species (genus Labroides) in the Indo-Pacific that specialize entirely on cleaning and have a characteristic color scheme – blue with a longitudinal black stripe – that allows their ‘clients’ to identify them as cleaners. Any such system can be exploited: and the saber-toothed blenny, Aspridonotus taeniatus, adopts the same color scheme as the labrids but when it gets close to the client it tears pieces of flesh out rather than picking off ectoparasites. In describing the various modes of fish feeding we have seen that the success of individuals at capturing prey is a consequence of having the right morphology and behavior. It is also important to realize that many species are not confined to just one mode of feeding. The tactics adopted by a species can change with age or size, time of day, and geographical location. On a moment-to-moment basis the behavior adopted by a fish is critical in determining the diet taken and the energetic consequences of food intake. It is assumed in modern studies of foraging that behaviors have been molded by natural selection in the same way as has morphology.

Foraging Strategies Given the assumption that behavior can be molded by natural selection, it becomes possible to analyze behaviors from an economic viewpoint. If a fish behaves so as to maximize its lifetime fitness, then in the short term it will choose to do things that maximize short-term gains such as increased growth rate or egg production and to minimize costs such as energy consumed or risk of predation during foraging. It then becomes possible to ask what behavioral strategy will maximize short-term gain or, in the jargon of foraging theory, optimize the behavior. With this approach it has been possible to predict what the optimal foraging strategy is for a species selecting prey from an environment with particular characteristics. The way in which tunas behave while foraging near ocean fronts can be understood with the aid of optimality arguments. Tuna in the bluefin group (genus Thunnus) travel widely in search of prey. They have often been observed aggregating at ocean fronts where warm water is separated from cooler water by a narrow transitional zone. It is characteristic of these fronts that the productivity of tuna food


378


is highest on the cool side of the front. This may be because the cooler water has recently upwelled and has higher plant nutrient levels. The dilemma facing the foraging tuna is that it prefers to be in the warmer water from a thermoregulation point of view but its best feeding opportunity is in the cooler water. Unlike many smaller fish, tuna have some control over their core body temperature. The vascular system has a heat exchange process by which blood moving from the center of the body outward passes vessels taking blood from the outside in. In this way, the core temperature of a tuna can be maintained significantly above ambient and controlled at a relatively constant level. For the tuna, this regulation becomes harder and more energetically costly in cool water, so that a prolonged stay in cool water could lead to death. The question for the tuna then is to decide how long it should stay foraging in the cool water where food availability is higher than in the more ‘comfortable’ warm water on the other side of the front. Using optimality methods borrowed from engineering, it is possible to model the physiology and behavior of the fish and to calculate the energetic costs and benefits of the fish being in either the undesirable cold water with high food or the desirable warm water with low food. The model predicts that the fish will behave optimally, that is, maximize its net energy gain, if it spends all of its nonfeeding time in the warm water, making quick sorties into the cold water area to fill its stomach. As soon as this has been achieved the fish withdraws again to the warm water to digest its meal. How long it has to stay in the cool water is a function of the abundance of food and the clarity of the water. The tuna is a visual predator, so the encounter rate with prey (prey met per unit time) will be a function of these two variables. Adopting this strategy will lead to the fish hovering around the boundary and, when applied to a school of tuna, may provide a mechanism for the observed aggregation behavior. For many species, food acquisition takes place in a competitive environment. As already mentioned, some species shoal together in an attempt to reduce the individual risk of predation. One cost of this behavior is that all the individuals in the group will be searching in the same area for the same type of food, although group foraging often means that food is found faster. The optimal behavior for an individual will then depend on what others choose to do. Individuals cope with this type of competition in a number of different ways. Experiments with groups of cod (Gadus morhua) in large aquaria show that access to food items delivered one at a time is determined largely by the visual acuity, swimming

speed, and hunger of each fish. The individuals that take the first few prey that are offered tend to be bigger than the others and hungrier, and may have a genetically determined higher basic metabolic rate. This type of competition by cod is usually termed ‘scramble competition’. Other species handle group competition in different ways. For example, the omnivorous damsel fish, Eupomacentrus planifrons, defends a territory against conspecifics, so ensuring for itself a private supply of food. A further method of coping with intraspecific competition is to develop a hierarchy so that individuals can recognize the status of others from behavioral signals. When confronted with a dominant, a subdominant will give way without a fight. In this way, the cost of contests is reduced, although the subdominant might be forced to feed as an opportunistic forager while the dominant takes a more selective diet. However, in the context of foraging in a group, the subdominant is doing the best it can. If competing individuals are genetically related, or live together for a long time, individuals might be prepared to give way to a competitor in any particular interaction over food. In this way, familiar or related competitors might operate on a tit-for-tat basis, thus sharing the resource. There is some evidence from three-spined sticklebacks (Gasterosteus aculeatus) that this occurs. Fish that have been living together spend less time chasing a partner that has caught first a prey offered simultaneously to them both than do fish that have met for the first time in the competitive arena. This outcome implies that individual fish recognize each other and can remember who did what when. Later work has shown that sticklebacks differentiate between familiar and unfamiliar partners using smell. Vision alone is not enough and it is likely that fish cannot recognize each other as known individuals. Situations in which the optimal behavior depends on how others behave are best handled theoretically using aspects of game theory. This predicts how rational decision makers should behave to maximize their payoffs in the face of competition. In fish behavioral studies, aspects of game theory have been used to predict how groups of sticklebacks should divide themselves when exploiting patches of food with different profitabilities and how individuals of the same species should behave when two or more are approaching a predator to undertake what is called ‘predator inspection’. Here individual prey fish suddenly leave their shoal and swim deliberately toward a predator before turning back and rushing back to the safety of the shoal. Such individuals are often accompanied by one or more conspecifics that



lag behind the leader. This behavior has been modeled as a cooperative interaction between the inspectors using a branch of game theory called the ‘prisoners’ dilemma’.

379

the seabird populations that have suffered a number of years with little or no fledging of young.

Feeding Behavior and Climate Change Food Chains Everything that has been said so far emphasizes links between fish at various levels in the ecosystem, as shown in Figure 1. A similar diagram could be drawn for any species of fish, meaning that the dynamics of marine ecosystems is a function of the relationships established through feeding. Certain fish species have key roles to play in that they are prey for a wide range of species. One such species in the North Sea is the sand eel (Ammodytes marinus), which is a major food item for herring, mackerel, cod, whiting, pollack (Pollachius pollachius), saithe (Pollachius virens), haddock, bass (Dicentrarchus labrax), turbot (Scophthalmus maximus), brill (Stemmatodus rhombus), megrim, plaice (Pleuronectes platessa), halibut (Hippoglossus hippoglossus), and sole. In addition, the sand eel is an important food item for many seabirds, particularly during the nesting season when, for example, the survival of puffin chicks (Fratercula arctica) depends on their parents bringing sufficient numbers of sand eels back to the nest. This one example shows how critical the links between species are in a marine ecosystem. Early attempts at fisheries management in the North Sea, and in most other areas of the world, ignored the interconnectedness of species through trophic interactions. Since the late 1980s, there has been an effort to take note of the interactions when fish stock assessment is carried out. One of the major effects of sustained fishing pressure on marine ecosystems has been the gradual reduction of abundance of the larger fish within a species and of the larger species. This has had the consequence of reducing the predation pressure on lower levels of the trophic web, so that species that have traditionally been given a low commercial value have increased in abundance and are all that is available. The sand eel illustrates this well. Until the early 1970s, there was no significant fishery for sand eels in the North Sea. The growing demands for fish meal, generated by the poultry and pig production industries, created a market for previously unused species such as sand eels. Coupled with reduced catches of higher-valued species such as herring and cod, this stimulated fishermen to focus on sand eels and this, together with the continued sustained high levels of effort on the predators of sand eels, is hastening the demise of the whole system. There have also been serious consequences for

In the first section it was made clear that many marine fish species start life as microscopic individuals that grow to be several orders of magnitude larger in length and weight. During the animal’s life it occupies a series of feeding niches, and this process has come to be known as an otogenetic niche shift. The process has important consequences for the ecology of fish. An ontogenetic niche shift is well illustrated by the North Atlantic cod where the eggs are numerous and small, larvae only a few millimeters long, and a 20-year-old adult can be over 1 m long. During its life history, the cod is most vulnerable when at its smallest size. This is partly because the larvae are prey to many planktonic organisms, but also because the survival and growth of the fish are dependent on them finding the right kind of food at the right time. In the North Sea, cod larvae appear in the plankton in the spring and one of their favourite foods is the copepod C. finmarchicus, which is also most abundant in the first part of the year. Over the past 20 years, data from the Continuous Plankton Recorder Survey, which has been mapping plankton distributions in the Northeast Atlantic for the past 70 years, has shown that C. finmarchicus abundance has fallen. C. finmarchicus favors cool water and is most abundant in northerly regions. Its place has been taken in the North Sea by Calanus helgolandicus, which is a warmer-water species and not so favored as food by cod. The timing of this species’ highest abundance is also not favorable to cod larvae as it is most abundant in the second part of the summer season. By this time, cod larvae have dropped out of the plankton and have become juveniles in inshore nursery areas. The change in plankton composition is thought to be due to the warming of the surface waters of the North Sea brought about by global climate change. As a result of this change in the plankton community, cod larvae are obtaining less food and dying in greater numbers, thus reducing the recruitment of new young fish to the adult population. This, coupled with continued overfishing, has brought the North Sea cod stock to a perilous state where it is almost at the point of commercial extinction. This example illustrates how critical feeding is to the well-being of a fish population. It also illustrates the complex interactions that link what happens to


380


individuals, large-scale events such as climate change, and events at the population level.

See also Benthic Organisms Overview. Coral Reef Fishes. Fisheries Overview. Fishery Management. Gelatinous Zooplankton. Habitat Modification. Large Marine Ecosystems. Mesopelagic Fishes. Pelagic Fishes. Plankton. Seabird Foraging Ecology. Upwelling Ecosystems.

Further Reading Beaugrand G, Brander KM, Lindley JA, Souisi S, and Reid PC (2003) Plankton effect on cod recruitment in the North Sea. Nature 426: 661--664. Brill RW (1994) A review of temperature and oxygen tolerance studies of tunas pertinent to fisheries oceanography, movement models and stock assessments. Fisheries Oceanography 3: 204--216. Gerking SD (1994) Feeding Ecology of Fish. San Diego, CA: Academic Press. Giraldeau L-A and Caraco T (2000) Social Foraging Theory. Princeton, NJ: Princeton University Press. Hardy AC (1924) The herring in relation to its animate environment. Part I: The food and feeding habits of the

herring. Fishery Investigations, London, Series II 7(3): 53. Hart PJB (1993) Teleost foraging: Facts and theories. In: Pitcher TJ (ed.) Behavior of Teleost Fishes, ch. 8. London: Chapman and Hall. Hart PJB (1997) Foraging tactics. In: Godin J-G (ed.) Behavioral Ecology of Teleost Fishes, ch. 5. Oxford, UK: Oxford University Press. Hart PJB (1998) Enlarging the shadow of the future: Avoiding conflict and conserving fish. In: Pitcher TJ, Hart PJB, and Pauly D (eds.) Reinventing Fisheries Management, ch. 17. Dordrecht: Kluwer. Hart PJB and Reynolds JD (eds.) (2002) Handbook of Fish Biology and Fisheries, Vol. I: Fish Biology, chs. 11–14 and 16. Oxford, UK: Blackwell. Helfman GS, Collette BB, and Facey DE (1997) The Diversity of Fishes. Oxford, UK: Blackwell. Lowe-McConnell RH (1987) Ecological Studies in Tropical Fish Communities. Cambridge, MA: Cambridge University Press. Pauly D, Christensen V, Dalsgaard J, Froese R, and Torres F, Jr. (1998) Fishing down marine food webs. Science 279: 860--863. Stephens DW and Krebs JR (1986) Foraging Theory. Princeton, NJ: Princeton University Press. Wootton RJ (1998) The Ecology of Teleost Fishes. Dordrecht: Kluwer.


FISH LARVAE E. D. Houde, University of Maryland, Solomons, MD, USA

in determining recruitment success can vary annually and seasonally.

Copyright & 2001 Elsevier Ltd. This article is reproduced from the 1st edition of Encyclopedia of Ocean Sciences, volume 2, pp 928–938, & 2001, Elsevier Ltd.

Introduction Most marine teleost (bony) fishes produce thousands to millions of planktonic eggs and larvae. Newly hatched larvae, usually 1–5 mm in length, are delicate, poorly developed, and retain many embryonic characteristics. They usually hatch with undeveloped mouth parts, fins, and eyes. Larvae drift and disperse in the sea. Most die before transforming into the juvenile stage. The larval stage originates as a nonfeeding, yolk-sac larva that derives nutrition from stored yolk and develops into an actively feeding larva that eventually transforms to a juvenile morphologically resembling a small adult. Transformation, which occurs from a few days to more than a year after hatching, often involves major changes in morphology (e.g. eels, flounders, herring) or may be less dramatic (e.g. cods, basses, sea breams). Lengths at metamorphosis are usually o25 mm, but can range from a few millimeters to many centimeters (some eels). The diverse, sometimes bizarre, suite of larval types that are collected represents a range of adaptations that promote survival and fitness in marine environments ranging from estuaries to the deep sea (Figure 1). Fisheries scientists and managers are concerned about growth and survival during the earliest life stages of fishes because variability in those processes can lead to 10-fold or greater differences in numbers of recruits that survive to catchable size. Causes of mortality are seldom evaluated. Since early in the twentieth century, scientists have realized that ocean circulation, frontal systems, and turbulence might be key physical factors controlling larval survival. Biological factors, especially larval nutrition and predation also are major controllers of survival and growth. The physical and biological factors combine to determine success of the reproductive effort, termed ‘recruitment’ by fisheries scientists. Recruitment processes are not confined to the larval stage but act on earlier (eggs) and later (juveniles) stages. The larval stage is important but does not stand alone as a ‘critical stage’, and its relative importance

Fish Larvae and Plankton Larvae of marine fishes, termed ichthyoplankton, usually are pelagic, drifting in the sea and interacting with pelagic predators and planktonic prey. Most fish larvae, even of species that ultimately are herbivores as juveniles or adults, are primarily carnivorous during the larval stage, feeding upon smaller planktonic organisms. In turn, larval fishes are the prey of larger nektonic and planktonic organisms. Escape from the precarious larval stage is accomplished via growth and ontogeny. Only a few individuals from thousands of newly hatched larvae survive the ever-present threats of starvation and predation during planktonic life. Eggs and larvae of marine fishes are collected in fine-meshed plankton nets or specially designed traps. Surveys at sea estimate distributions, abundance, diversity, and structure of ‘ichthyoplankton’ communities, including associations of larvae with their predators and prey. Such surveys sometimes are a component of stock assessments used in fisheries management. Larval distributions within nursery areas, including vertical distributions, differ among species. Most larvae are in the upper 200 m of the water column, although buoyant eggs, which gradually rise towards surface after spawning, may be deeper if adults spawned at depth. Accurate estimates of abundances and production of eggs and small larvae are possible because these stages usually cannot avoid samplers. Declines in abundances of older (and larger) larvae in plankton collections provide the means to estimate mortality rates if dispersal losses by water currents can be determined, and if the probability of sampler avoidance is known.

Larval Survival The larval stage may be most important in controlling levels of subsequent recruitment. Early in the twentieth century, Johann Hjort (1914) offered the ‘critical period’ hypothesis, proposing that incidence of starvation at the time of yolk-sac exhaustion, when larvae first required plankton as food, was the primary factor determining variability in year-class recruitment success. In combination, probabilities of starvation and transport to unsuitable nursery


381

382

FISH LARVAE

habitats constitute the two elements of Hjort’s hypothesis. Although the hypothesis cannot be rejected after 85 years, it is now apparent that recruitment variability is generated by a multitude of processes acting throughout the early-life stages in fishes. Hjort’s hypothesis provided a foundation for subsequent attempts to explain recruitment variability. The potential for high and variable larval stage

mortality led many scientists to hypothesize that coarse controls on the magnitude of recruitment are set during the larval stage rather than later in life. The ‘match-mismatch’ hypothesis proposed by Cushing builds upon Hjort’s ideas, emphasizing the importance of temporal coincidence in spawning and bloom dynamics of plankton, the primary food of larval fish. There is support for this hypothesis,

1

3

2

4

5 6

8 7

10

11

9

12

14 13


FISH LARVAE

especially for species with short spawning seasons in high-latitude seas. A ‘match’ between spawning and spring plankton blooms ensures larval growth and survival, while a ‘mismatch’ results in high mortality. The ‘stable ocean’ hypothesis, proposed by Lasker in the 1970s, is another nutrition-related explanation for variability in larval survival. Lasker hypothesized that relaxation of storm winds and intense upwelling resulted in a stable, vertically stratified ocean in which strata of fish larvae and their prey coincide, promoting larval nutrition and survival. The hypothesis is strongly supported for species inhabiting coastal upwelling regimes. Contrasting with the match-mismatch hypothesis is the proposal by Sinclair and Iles that gyre-like circulation features in well-mixed coastal waters define spawning areas while retaining eggs and larvae. This hypothesis emphasizes physics and circulation features, rather than nutritional factors, as the controller of recruitment variability. The hypothesis gains support from evidence on recruitment of numerous herring (Clupea harengus) and some cod (Gadus morhua) stocks in the North Atlantic. Sinclair and Iles argued that high larval retention promotes strong recruitment while failed retention diminishes success. Retention not only reduces larval losses to dispersal, but ensures genetic integrity of a stock by confining reproduction within the retention area. Although this hypothesis differs from the ‘match-mismatch’ hypothesis, there is evidence from

383

some cod stocks that ‘match-mismatch’ and retention mechanisms operate together to promote larval survival. In the tropics and especially on coral reefs, scientists have debated whether supply of larvae or postsettlement mortality of juveniles is the primary factor generating variability in recruitment levels. The ‘lottery’ hypothesis, of Sale, proposes that fish larvae are delivered by ocean currents to reefs where they settle onto structure that, if by chance is free of predators or competing fish, will lead to successful recruitment. Recent evidence supports both larval supply and postsettlement controls as important for recruitment success. Also, there is accumulating evidence that longrange dispersal, once assumed to dominate early-life processes in tropical seas, may in fact be limited and that retention near islands and reefs is common as larvae develop from hatching to settlement stages.

Foods and Feeding Fish larvae must feed frequently to ensure fast growth, a prerequisite for high survival. Successful initiation of feeding depends upon availability of suitable kinds and sizes of prey, which usually is zooplankton 50–100 mm in width. Sizes of prey that are consumed are strongly related to larval size (specifically mouth size). Nauplii of copepods are perhaps the most common prey of small fish larvae.

Figure 1 (Left) Larval fishes. (1) Sardinella zunasi (Clupeidae), 4.8 mm. (Reproduced with permission from Takita T (1966).) Egg development and larval stages of the small clupeoid fish Harengula zunasi Bleeker and some information about the spawning and nursery in Ariake Sound. Bulletin of the Faculty of Fisheries, Nagasaki University 21: 171–179. (2) Clupea pallasi (Clupeidae), 10.4 mm. (Reproduced with permission from Matarese AC, Kendall AW Jr, Blood DM and Vinter BM (1989).) Laboratory Guide to Early Life History Stages of Northeast Pacific Fishes. NOAA Technical Report, NMFS 80 Seattle, WA: National Marine Fisheries Service.) (3) Diaphus theta (Myctophidae), 4.6 mm. (Reproduced with permission from Matarese et al., 1989.) (4) Benthosema fibulatum (Myctophidae), 8.7 mm. (Reproduced with permission from Moser HG and Ahlstrom EH (1974).) Role of larval stages in systematic investigations of marine teleosts: the Myctophidae, a case study. Fishery Bulletin, US 72: 391–413. (5) Sebastes melanops (Scorpaenidae), 10.6 mm. (Reproduced with permission from Matarese et al., 1989.) (6, 7) Theragra chalcogramma (Gadidae), 6.2 and 13.7 mm. (Reproduced with permission from Matarese et al., 1989.) (8) Carangoides sp. (Carangidae), 4.4 mm. (Reproduced with permission from Leis JM and Trnski T (1989).) The Larvae of Indo-Pacific Shorefishes. Sydney: University of New South Wales Press.) (9) Plectranthrias garupellus (Serranidae), 5.5 mm. (Reproduced with permission from Kendall AW Jr (1979).) Morphological Comparisons of North American Sea Bass Larvae (Pisces: Serranidae). NOAA Technical Report NMFS Circular 428. Rockville, MD: National Marine Fisheries Service. (10) Lutjanus campechanus (Lutjanidae), 7.3 mm. (Reproduced with permission from Collins LA, Finucane JH and Barger LE (1980).) Description of larval and juvenile red snapper, Lutjanus campechanus. Fishery Bulletin, US 77: 965–974. (11) Naso unicornis (Acanthuridae), 5.9 mm. (Reproduced with permission from Leis JM and Richards WJ (1984).) Acanthuroidei; development and relationships. In: Moser HG, Richards WJ, Cohen DM, Fahay MP, Kendall AW Jr, and Richardson SL (eds) Ontogeny and Systematics of Fishes, pp. 547–551. American Society for Ichthyologists and Herpetologists, Special Publication 1. (12) Xiphias gladius (Xiphiidae), 6.1 mm. (Reproduced with permission from Collette BBT, Potthoff T, Richards WJ, Ueyanagi S, Russo JL and Nishikawa Y (1984)) Scombroidei: development and relationships. In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, pp. 591–620. American Society of Ichthyologists and Herpetologists, Special Publication 1. (13) Thunnus thynnus (Scombridae), 6.0 mm. (Reproduced with permission from Collette BB et al., 1984.) In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, 591–620. American Society of Ichthyologists and Herpetologists Special Publication 1.) (14) Paralichthys californicus (Paralichthyidae), 7.0 mm. (Reproduced with permission from Ahlstrom EH, Amaoka K, Hensley DA, Moser HG and Sumida BY (1984).) In: Moser HG et al. (eds) Ontogeny and Systematics of Fishes, pp. 640–670. American Society of Ichthyologists and Herpetologists, Special Publication 1.)


384

FISH LARVAE

Concentrations of nauplii and other zooplankton have often been analyzed to evaluate feeding conditions for larvae in the sea. Recent research has demonstrated that many species of fish larvae initiate feeding on a diverse spectrum of small planktonic organisms, including some phytoplankton and protozoa that are often more abundant than copepod nauplii. Laboratory experiments indicate that concentrations of copepod nauplii exceeding 100 per liter may be required to support feeding and growth of larval fish, leading to hypotheses and simulation models that implicate patchiness of prey and smallscale turbulence as mechanisms promoting prey encounter, successful feeding, and growth. Such enhancing mechanisms clearly are important, but if diverse diets are the norm, as recent evidence suggests, then availability of suitable larval prey in the sea may be higher than once believed. As development and growth proceed, feeding success increases. Maximum prey sizes in larval diets increase substantially, but the mean size increases only slowly because larvae continue to consume small prey while adding larger prey to the diet. This ontogenetic shift occurs gradually in most fishes, but in some mackerels, tunas, and other species the shift is dramatic, and piscivory on larval fish becomes their primary source of nutrition during the earliest larval stage. Large amounts of prey must be consumed.

Marine fish larvae typically consume >50% of their body weight daily to achieve average growth. Some larvae from warm seas, e.g. tunas and anchovies, consume >100% of their body weight each day to grow at average rates. Such high food requirements have reinforced the belief that poor feeding conditions, slow growth, and starvation are major causes of larval mortality in the sea. Starving fish larvae are seldom observed in the sea because such larvae are believed to be selectively preyed upon, then disintegrating rapidly in stomachs of predators. Thus, dead or starving larvae are infrequent in ichthyoplankton collections. Nutritional condition of fish larvae can be evaluated by morphological, histological, and biochemical approaches. Nucleic acid analysis, when properly conducted, shows that poorly fed larvae in environments with low prey have low RNA/ DNA ratios, indicating poor potential for growth. Fish larvae that fail to initiate feeding soon after yolk depletion will become nutritionally compromised and poorly conditioned, reaching a ‘point-of-no-return’ within 2–10 days, from which they cannot recover even if adequate food is provided. Little is known about specific feeding behaviors of fish larvae. Larvae mostly feed visually and must be in close proximity (e.g. less than one body length) to successfully encounter and capture prey. Most feeding occurs during daylight, although threshold light levels

(A)

(B) Figure 2 Typical feeding behavior of an Atlantic herring larva. (A) ‘Tube’ search. Only those food items in the tube are available to the larva. (B) S-flex feeding. Sequence of stages during an unsuccessful feeding attempt by a larva. (Reproduced with permission from Rosenthal H and Hempel G (1970).) Experimental studies in feeding and food requirements of herring larvae (Clupea harengus L.) In: Steele JH (ed.) Marine food chains, pp. 344–364. Berkeley: University of California Press.)


FISH LARVAE

are low, in the range 0.01–0.10 lux, levels equivalent to dawn and dusk periods. In elongate, herring-like larvae, a typical ‘S-flex’ behavior has been described in which a larva, upon encountering a prey, flexes its body into an S-shape before striking at the prey (Figure 2). Other kinds of larvae use a modified ‘S-flex’ behavior or a ‘cruise and strike’ searching behavior. As in other life stages, successful feeding by larvae depends on the relationship: P ¼ E A C, where E, A, and C are probabilities of encountering, attacking, and capturing a prey. Modeling and understanding this relationship, and changes in it during larval development, are important to evaluate larval survival potential.

Predation Much evidence indicates that predation is the major direct cause of mortality to early-life stages of fish. This conclusion does not exclude starvation and nutritional deficiencies as causes of, or contributors to, larval mortality. Because predation in the sea is strongly size-dependent and predators are size-selective, fish larvae that are slow-growing or starving will remain in small, highly vulnerable size classes longer and may have higher predation risk. Thus, larval growth rate and its dependence on nutrition links the predation and starvation processes. Larval growth rate under many circumstances is an important measure of susceptibility to predation. There are many kinds of predators on fish eggs and larvae. Predators include jellyfishes (ctenophores and medusae), fish, and crustacea (e.g. euphausiids, large copepods). Cannibalism on eggs and larvae occurs and can be an important population regulatory mechanism. The relative importance of the various predators is poorly known. Although a suite of predators of different taxa and sizes eat larvae of preferred sizes, the integrated effects of complex, multispecific predation on recruitment outcomes have hardly been evaluated in the sea or in laboratory experiments. Models of predation also generally have focused on single predators of near-uniform size preying on fish larvae of a single species. In natural ecosystems a community of predators of varying kinds and sizes, which are distributed patchily in time and space, interacts with an assemblage of fish early-life stages. Outcomes are difficult to predict and will be modified by availability of alternative prey resources. Vulnerability of larval fish to a predator of defined size can often be represented by dome-shaped responses in which vulnerability peaks at an intermediate larval size or by a consistent decline in

385

vulnerability as larval size increases (Figure 3). Predation rate on larvae is frequently a dome-shaped function of the ratio of prey size : predator size, where larval fish (the prey), are consumed most efficiently when they are 5–15% of a predator’s length, a ratio that is consistent for predators as diverse as jellyfish or fish. Vulnerability of larval fish to predation is the product of encounter rate (dependent upon abundances, sizes, and swimming speeds of both predators and prey) and susceptibility (attack probability capture probability) of larvae. Thus, sizes and swimming speeds of predators and larvae determine vulnerability of an individual larva. At the population level, size-specific abundances of the predator clearly will be important in determining whether a particular predator can control or inflict significant mortality on a population of larval fish. Many common predators, e.g. some medusae and fish, span a broad range of sizes and may be predators on populations of fish larvae over a wide size range. Although size is the dominant factor determining vulnerability of fish larvae to predation, ontogeny plays an important role. Fish larvae can react to predators by sensing vibrations via free neuromasts that are present at hatching. As growth occurs, neuromasts proliferate, the lateral line develops, and vision may become increasingly important to detect and escape predators. Development of musculature and fins, especially the caudal fin complex, increases the swimming power and escape ability of larval fishes. Ontogenetic improvements in larval ability to avoid predators are counteracted to an extent by the increased sizes of predators targeting larger larvae and juvenile fish. Mortality rates of large larvae and juveniles generally decline, indicating that survival advantages attained through ontogeny and growth outweigh increased suitability of these fish to a suite of larger predators.

Temperature and Salinity Temperature exerts strong control over rate processes associated with metabolism and growth in larval fishes. Temperature and body size may be the two most important controlling variables in the early lives of fishes. Temperature also controls rate processes in predators and thus indirectly controls rates at which larvae die from predation. Furthermore, seasonal development of planktonic prey of fish larvae depends in part upon temperature, and is especially important in seasonally variable mid and high latitudes where combined effects of temperature and prey levels control larval growth potential. Salinity


386

FISH LARVAE

Susceptibility

Encounter rate

Vulnerability

Ambush raptorial invertebrate (A)

Cruising invertebrate

Relative scale

(B)

Filter-feeding fishes (C)

Raptorial fishes (D)

Relative larval size Figure 3 Conceptual models showing relative encounter rates, susceptibility, and vulnerability of fish larvae to different predator types. (Reproduced with permission from Bailey KM and Houde ED, 1989.)

generally is of secondary importance, but it too can be critical for some anadromous fishes whose larvae inhabit estuarine zones along the salinity gradient or at the interface between fresh water and salt waters. Growth rates of marine fish larvae are surprisingly plastic. Weight-specific growth rates may vary by more than fourfold in response to temperature differences within tolerable ranges, which imparts high variability

to larval stage durations. Although temperature can exercise physiological control over growth, prevailing temperatures in many continental shelf and oceanic nursery areas differ only slightly during a spawning season or between years, and larval growth may in fact be controlled more by prey availability in those ecosystems. However, in shallow estuaries and neritic habitats, fluctuations in temperature induced by


FISH LARVAE

weather fronts or shifts in circulation patterns may be sufficiently strong to influence larval growth rates or to directly cause mortality. Ranges of temperature tolerated by larvae are generally lower than for juveniles and adults. Temperature and body size control bioenergetics relationships in marine poikilotherms, including larval fish. Growth rates, metabolic rates, and possibly growth efficiencies and assimilation rates of fish larvae, are sensitive to changes in temperature. Each fish species has a temperature at which it performs optimally. In meta-analyses across species and ecosystems, larval growth rates were demonstrated to increase as temperature increased. Species from high latitudes grow slowly (often o10% body weight per day) while tropical species grow fast (>30% per day). In the meta-analysis, mortality rates are positively correlated with growth rates, and thus also are temperature-dependent. Salinity controls water balance in osmotic regulation. Marine fish larvae often have surprisingly broad tolerances of salinity. Even anadromous fishes, whose larvae normally live in salinities o1 psu, may perform well in laboratory experiments at salinities of 5–10 psu. At high salinities marine fish larvae drink sea water to regulate water balance. Drinking rates increase as either salinity or temperature increases. Salinity and temperature often act together to affect the physiological performance of larvae. In particular, optimum salinity–temperature combinations determine hatching success and larval performance. Moreover, temperature and salinity combinations in the sea determine the density field ðst Þ in which larvae are located. The density field controls pycnocline and thermocline depths, which define strata, discontinuities, shear zones, and transition depths that may aggregate prey and predators of fish larvae as well as the larvae themselves.

Behavior Except for larval feeding behavior (see above), relatively little is known about the behavior by individuals with respect to environmental cues or other stimuli. Swimming behaviors are documented for some species. Larvae with elongate bodies (e.g. sardines) swim with an anguilliform motion, while larvae with more compact bodies (e.g. cod-like or basslike larvae) adopt a modified carangiform, ‘cruisepause’ swimming behavior. When searching for food, swimming typically is slow, one or two body lengths per second, and is the predominant behavior of feeding-stage larvae during daylight hours. Feedingstage larvae generally can cruise at sustained

387

swimming speeds of 3–10 body lengths per second, but newly-hatched, yolk-sac larvae are incapable of strong, directed swimming in a horizontal plane. Burst swimming to escape predators, at >10 body lengths per second is observed. Predator detection and avoidance behaviors become increasingly effective during ontogeny as sensory systems develop. Fish larvae may migrate vertically over tens of meters on a diel basis. This ability allows larvae to track food by adjusting depth distributions to coincide with depths where prey is abundant. In addition, larvae can potentially avoid parts of the water column where predators prevail or where environmental conditions and water quality (e.g. low dissolved oxygen) are unfavorable. Dispersal of larvae or, alternatively, retention on nursery grounds can depend on depth selection that promotes ‘selective tidal stream transport.’ This behavioral mechanism requires that larvae coordinate vertical migrations with tides and currents to ensure transport to, or retention within, favorable nursery areas. This mechanism is particularly important in estuaries, but may also regulate larval drift or retention on continental shelf nursery areas.

Integration Larval populations of many marine fishes are distributed over tens to thousands of kilometers and have early-life durations that range from days to months. However, the fates of individual larvae depend upon interactions with their environment, prey, and predators measured on spatial scales of micrometers to meters and on timescales of fractions of seconds to hours. Events at all spatial and temporal scales are potentially important and can generate significant variability in larval survival. The potential for recruitment success is indicated by both survival and proliferation of cohort biomass during early life. While numbers decline steadily, biomass of successful cohorts increases. Stage-specific survival during early life depends upon relative rates of mortality (M) that reduce a cohort’s numbers and growth (G) that controls accumulation of individual mass. Biomasses of most marine fish cohorts decline during the larval stage to levels o1% of their biomasses at hatching because M/G41.0. In many marine and estuarine fish larvae (e.g. walleye pollock Theragra chalcogramma, American shad Alosa sapidissima, bay anchovy Anchoa mitchilli), the body size at which the transition from M/G41.0 to o1.0 occurs, after which cohort biomass increases, varies annually. Cohorts that are strong contributors to recruitment make the transition at smaller size than unsuccessful cohorts (Figure 4).


388

FISH LARVAE

M/G minimum B M= G M >G Biomass decreasing

M

Encyclopedia of Ocean Sciences, Second Edition, Volume 2: D-F

Encyclopedia of Ocean Sciences, Second Edition, Volume 3: G-M

Encyclopedia of Ocean Sciences, Second Edition, Volume 1-6

Encyclopedia of Ocean Sciences, Second Edition, Volume 1: A-C

Encyclopedia of Ocean Sciences, Second Edition, Volume 5: S

Encyclopedia of Ocean Sciences, Second Edition, Volume 6: T-Z